# Tectonic structures and stress state associated with seamount subduction in accretionary prism.

## Implications for slip behavior

Alexis Gauthier, Nadaya Cubas, Laetitia Le Pourhiet

► To cite this version:

Alexis Gauthier, Nadaya Cubas, Laetitia Le Pourhiet. Tectonic structures and stress state associated with seamount subduction in accretionary prism. Implications for slip behavior. Earth and Planetary Science Letters, 2025, 675, pp.119767. 10.1016/j.epsl.2025.119767. hal-05477578

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# Tectonic structures and stress state associated with seamount subduction in accretionary prism. Implications for slip behavior

Alexis Gauthier<sup>a, *, ID</sup>, Nadaya Cubas<sup>a</sup>, Laetitia Le Pourhiet<sup>a, b ID</sup>

<figure><img src="image_3.png" /></figure>

## ARTICLE INFO

Editor: Dr C Carolina Lithgow-Bertelloni

### Keywords:

- Seamounts
- Accretionary prism
- Subduction
- Gravitational collapse
- Tectonic overpressure
- Seamount decapitation
- Shallow Slow Slip events

## ABSTRACT

Seamounts on oceanic plates are thought to affect the dynamics of accretionary prism at subduction zones, particularly the slip behavior. Geophysical studies show that pore fluid overpressure is associated with the nucleation of shallow slow slip events (SSEs). We use numerical simulations in which the subducting plate bends in response to tectonic, gravitational collapse and surface processes to capture structural complexities arising from the subduction of buried seamounts and compute the resulting stress state. We show that tectonic slices of undeformed sediments develop on both sides of the seamount and that a basin forms atop the seaward slice. These structures, recognizable in seismic data, represent the long-term tectonic signature of seamount subduction.

Gravitational collapse only occurs when a critical lengthscale, we quantified, of perturbed accretionary prism is reached. Seamounts that are small enough, instead get decapitated and do not perturb the prism over a sufficient lengthscale to trigger gravitational collapse.

Stress analysis reveals that the seaward side of the seamount is under an extensional regime, whereas its landward side is in a compressive state. We thus propose that sustained pore fluid overpressure can only develop on the landward side. This mechanism supports observations along Hikurangi, where a highly reflective zone, interpreted as a zone of pore fluid overpressure, has been imaged downdip of the Papaku seamount. Since this zone coincides with the largest slip magnitudes of shallow SSEs, we propose that the landward slice of undeformed sediments presents favorable conditions for SSEs nucleation.

## 1. introduction

Seamounts are ubiquitous features on oceanic plates and, consequently, in subduction zones (Hillier and Watts, 2007; Wessel et al., 2010). The impact of a subducting seamount on the overriding plate has been extensively investigated through both analogue and numerical modeling (Dominguez et al., 1998, 2000; Ruh, 2016; Noda et al., 2020; Wang et al., 2021; Miyakawa et al., 2022; Pajang et al., 2022a) These studies, along with observational data, have shown that seamounts can significantly damage the upper plate (Dominguez et al., 1998, 2000; Marcaillou et al., 2016), reducing its ability to store elastic energy compared to an undisturbed upper plate (Wang and Bilek, 2011, 2014). Rough subduction interfaces are therefore associated with aseismic creep (Wang and Bilek, 2014), and seamounts are thought to act as barriers to earthquake rupture propagation (Kodaira et al., 2000; Duan, 2012; Geersen et al., 2015; Cubas et al., 2022). These hypotheses are

based on the idea that seamounts are very rarely decapitated which has been challenged by recent observations along the Franciscan Complex (Bonnet et al., 2024).

Along the dip of the interface, seamounts are also suggested to promote shallow slow slip events (SSEs), as observed along the Hikurangi, Nankai, and Ecuador margins (Kodaira et al., 2004; Bell et al., 2010, 2014; Yokota et al., 2016; Collot et al., 2017; Barker et al., 2018; Todd et al., 2018; Shiraishi et al., 2020). SSEs are commonly associated with high pore fluid pressure, as inferred from geophysical surveys (Bell et al., 2010; Audet et al., 2010; Saffer and Tobin, 2011; Saffer and Wallace, 2015; Barker et al., 2018; Flores et al., 2025) and supported by numerical modeling based on rate-and-state friction laws (Liu and Rice, 2007, 2009).

However, unlike deeper SSEs, which are thought to occur within specific temperature windows related to dehydration reactions (Peacock, 2009; Audet et al., 2010), the high pore fluid pressure

<footnote><sup>a</sup> ISTE>P, Sorbonne Université, 4 Pl. Jussieu, 75005 Paris, France</footnote>

<footnote><sup>b</sup> Institut Universitaire de France, France</footnote>

<footnote>* Corresponding author.</footnote>

<footnote>E-mail address: alexis.gauthier@sorbonne-universite.fr (A. Gauthier).</footnote>
associated with shallow SSEs is suggested to be triggered by a compositionally and geometrically heterogeneous plate interface (Saffer and Wallace, 2015). Since long-term fluid overpressurization can be attributed to the persistent stress state around the seamount, this hypothesis can be further investigated using numerical approaches (Ellis et al., 2015; Sun et al., 2020; Miyakawa et al., 2022). However, capturing the long-term overpressurization requires consideration of the structural inheritance induced by seamount subduction. Although stress states have been analyzed using numerical approaches (Ellis et al., 2015; Ruh, 2016; Sun et al., 2020), inherited structures have not been systematically taken into account.

In addition, despite extensive investigations of seamounts through both analogue and numerical modeling, none of these studies have simultaneously accounted for low basal friction expected in subduction zones (Cubas et al., 2013; Gao and Wang, 2014), surface processes induced by an evolving topographic slope, realistic mechanical properties allowing for seamount deformation, and lithospheric flexure, all key factors controlling wedge structures (Davis et al., 1983; Konstantinovskaya and Malavieille, 2011; Ruh et al., 2012; Ruh, 2020; Pajang et al., 2022a).

Most of these previous studies have focused on isolated subaerial seamounts, whereas seamounts are often organized in chains. Moreover, seamounts located up-dip of SSEs along the Hikurangi and Nankai margins are buried beneath a thin sedimentary cover (Wallace, 2020; Lallemand et al., 2024).

Therefore, we here perform numerical simulations of accretionary prisms disrupted by multiple buried seamounts of varying sizes, incorporating realistic properties and boundary conditions in order to describe the evolving structural complexities induced by subducting seamounts and to evaluate the mechanical conditions for their decapitation.

Based on these results, we propose structural and morphological criteria to identify active and past buried seamounts. These criteria can be used in poorly imaged subduction zones to ultimately increase the number of documented cases. We test them against two seismic profiles from the Nankai and Hikurangi margins.

We also describe the heterogeneities in stress state induced by the structural complexities. By comparing the tectonic structures and the tectonic rock over- or under-pressure derived from our numerical simulations, we discuss their potential impact on the long-term pore fluid pressure — regions where shallow SSEs are thought to nucleate.

## 2. – method

### 2.1. – governing equation

pTatin2d (May et al., 2014, 2015) code employs the finite element method on a Lagrangian-Eulerian grid, combined with the material point method. Since only brittle deformation was considered in this study, the thermal part of the code was disregarded.

At each time step, the Stokes equations for velocity (v) and rock pressure (P) are solved on a Q2-P1 grid, ensuring the conservation of momentum (eq. (1)) for an incompressible fluid (eq. (2)).

$$ \nabla \cdot \eta(\nabla \mathbf{v} + (\nabla \mathbf{v})^t) - \nabla P + \rho g = 0, \quad (1) $$

$$ \nabla \cdot \mathbf{v} = 0. \quad (2) $$

The coefficients include the effective viscosity ($\eta$), gravity acceleration vector ($\mathbf{g}$), and material density ($\rho$).

To determine the effective viscosity, we use the Drucker-Prager yield criterion (eq. (3)):

$$ \eta = \frac{\sigma_{DP}}{2\epsilon} = \frac{C\cos\phi + P\sin\phi}{2\epsilon^{\cdot II}}, \quad (3) $$

where C represent the cohesion, $\phi$ the friction angle and $\epsilon^{\cdot II}$ the second invariant of the strain rate tensor.

Markers carry information about rocks, particularly their rheology. At each time step, the velocity is projected onto the markers which are then advected.

To simulate strain softening, friction linearly decreases from $\phi_{max}$ to $\phi_{min}$ as strain accumulates ($\epsilon_p$) from $\epsilon_{min} = 0$ to $\epsilon_{max} = 0.5$:

$$ \phi = \phi_{max} - \frac{\epsilon_p - \epsilon_{min}}{\epsilon_{max} - \epsilon_{min}} (\phi_{max} - \phi_{min}) \quad (4) $$

Similarly, cohesion (C) decreases from $C_{max}$ to $C_{min}$. Softening is applied to the sediments and oceanic crust, but not to the backstop or the décollement (Table 1). The backstop is not affected by weakening, as it is not expected to deform in this type of experiment. Similarly, no softening is applied to the décollement in order to maintain full control over its properties for the parametric study.

Finally, to account for surface processes, we use the Culling diffusive erosion law to smooth the topography (Jourdon et al., 2018):

$$ \frac{\partial h}{\partial t} = -\nabla \cdot (\kappa \nabla h), \quad (5) $$

with h the altitude and $\kappa$ the diffusion coefficient.

### 2.2. – flexure of the lithosphere

To account for the lithospheric flexural response to sediment loading, a new boundary condition was implemented to allow bending at the base of the model (Fig. 1).

At each time step, the surface velocity ($v_y|_{surface}$) and the duration ($dt$) are extracted to calculate the change in topography ($\Delta h$). An image of $\Delta h$ with respect to the ordinate axis is generated to perform its sine decomposition (sum of 100 sinus).

For each wavelength ($\lambda_i$), we calculate the flexural response ($\Delta\omega_o(\lambda_i)$) associated with a sinusoidal change in topography ($\Delta h(\lambda_i)$). We use an analytical solution from Turcotte & Schubert (2002), modified to account for the submersion of the prism:

$$ \Delta\omega_o(\lambda_i) = \frac{\Delta h(\lambda_i)}{\frac{\rho_m - \rho_w}{\rho_s - \rho_w} - 1 + \frac{D}{(\rho_s - \rho_w)g} \left(\frac{2\pi}{\lambda_i}\right)^4} \quad (6) $$

where $\rho_m$, $\rho_s$, and $\rho_w$ are the density of the mantle, the sediment, and the water, respectively. D is the rigidity of the lithosphere associated with the elastic thickness ($T_e$) (Turcotte and Schubert, 2002):

$$ D = ET_e^3 12(1 - \nu^2) \quad (7) $$

with E (=10 GPa) and $\nu$ (=0.25) the Young's modulus and Poisson's ratio respectively.

Finally, to determine the total flexural response, we sum the flexural responses associated with each wavelength. The total flexural response is then applied to the thermal and mechanical grids, and the markers are advected without considering the image part.

### 2.3. – model set up

The domain is discretized using a mesh of 700 × 16 Q2-P1 elements

<caption>Table 1</caption>

Rheological Parameters. C is the cohesion, $\phi$ the friction angle and $\rho$ the density.

<table><thead><tr><th></th><th>Sediments</th><th>Seamount</th><th>Backstop</th><th>Décollement</th></tr></thead><tbody><tr><td>C<sub>max</sub> - C<sub>min</sub> (MPa)</td><td>2-1</td><td>20-10</td><td>20</td><td>0.1</td></tr><tr><td>φ<sub>max</sub> - φ<sub>min</sub> (°)</td><td>20 - 10</td><td>30</td><td>30</td><td>2, 5 or 8</td></tr><tr><td>ρ</td><td>2.4</td><td>2.9</td><td>2.8</td><td>2.4</td></tr></tbody></table>
<figure><img src="image_4.png" /><figcaption>**Fig. 1.** Model Set Up.</figcaption></figure>

$v_x$ represents the velocity of the subducting plate. The top boundary behaves as a free surface. The weight of the water column is prescribed as stress normal to the deformed boundary below sea level. In yellow are represented the syn-tectonic sediments deposited before the initiation of the convergence, during the initiation phase.

over a horizontal length of 500 km. The mesh is refined along the y-axis to ensure a high resolution at the base of the model (Fig. 1). For computational efficiency, only sediments are explicitly included in the computational domain for the oceanic plate (Fig. 1). A simulation with an explicit oceanic crust is provided in Supplementary Material (M2_withcrust), showing that results are similar. The sediments are of 2 km thick and have sandstone/quartz properties (Table 1). They are depicted as alternating layers of light blues, which lack rheological significance but serve to visually track tectonic deformation as in analog modeling. At the bottom, a décollement layer of 200 m thick allows localization of the deformation at the crust-sediment boundary (dark blue on Fig. 1, Table 1). To investigate the effect of the décollement friction on the deformation, three different frictions were tested: 2, 5 and 8°. The décollement is interrupted by seamounts, with oceanic crust properties (Table 1), buried under the sediment cover. To investigate the effect of their size on the accretionary prism deformation, we varied their width and height (3 km wide/0.7 km height and 4 km wide/1.2 km

<caption>**Table 2**</caption>

Variable Model Parameters. EET is the elastic thickness of the lithosphere. The distance between the first seamount and the left boundary of the domain is 109 km by default, except for M2', where this distance is 107 km as an offset is applied.

<table><thead><tr><th></th><th>Seamount Size</th><th>Décollement Friction [°]</th><th>EET [km]</th><th>1st seamount offset</th></tr></thead><tbody><tr><td>M1</td><td>4 km wide and 1.2 km high</td><td>5</td><td>25</td><td>—</td></tr><tr><td>M2</td><td>3 km wide and 0.7 km high</td><td>5</td><td>25</td><td>—</td></tr><tr><td>M2'</td><td>3 km wide and 0.7 km high</td><td>5</td><td>25</td><td>Yes</td></tr><tr><td>sF1</td><td>4 km wide and 1.2 km high</td><td>2</td><td>25</td><td>—</td></tr><tr><td>IF1</td><td>4 km wide and 1.2 km high</td><td>8</td><td>25</td><td>—</td></tr><tr><td>sE1</td><td>4 km wide and 1.2 km high</td><td>5</td><td>10</td><td>—</td></tr><tr><td>IE1</td><td>4 km wide and 1.2 km high</td><td>5</td><td>40</td><td>—</td></tr><tr><td>sF2</td><td>3 km wide and 0.7 km high</td><td>2</td><td>25</td><td>—</td></tr><tr><td>IF2</td><td>3 km wide and 0.7 km high</td><td>8</td><td>25</td><td>—</td></tr></tbody></table>

An additional simulation, M2_with_crust, is provided in the Supplementary Material. It is similar to M2, but includes 2 km of explicitly represented oceanic crust.

height, Table 2). A backstop, shown in brown and with continental crust properties (Table 1) is placed on the left side of the model.

At the base of the model, a constant velocity ($v_x = 5$ cm.yr$^{-1}$) is imposed to represent the displacement of the subducting plate. The oceanic crust and mantle are implicitly represented, and hence considered as non-deformable. Sea level is set at 5.5 km above the sediment surface, on the incoming side of the domain, where the height is not affected by lithospheric flexure (Fig. 1). In the simulations, it is accounted for by a stress applied normal to the free surface that is equal to the hydrostatic pressure of the overlying water (Pajang et al., 2022b).

At the left boundary, sediments enter the model at the plate velocity ($v_x = 5$ cm.yr$^{-1}$). At the right boundary, a zero-horizontal velocity is imposed, except for the first 200 m at the base which allows for the accretion of sediments against the backstop and the exit of the décollement layer from the model domain (i.e. undergoes subduction). At the top boundary, a free surface is imposed to capture the topography. Before initiating compression, simulations are run with zero velocities for 1 Ma, to allow the topography to equilibrate according to the Culling diffusive erosion law. The sediments deposited during this phase, referred to as syn-tectonic sediments, share the same rheology as the initial sediments.

### 3. – results

#### 3.1. – reference simulation

The Fig. 2 illustrates the reference simulation M1 (Table 2) depicting the subduction of two seamounts divided into eight stages. At each stage, we show the tectonic structures, as well as the location and nature of the deformation. Deformation localization is visualized using the $\dot{e}_{xx}$ component of the strain rate tensor (S. Pajang et al., 2022a). As only brittle deformation is considered, shear bands exhibits frictional behavior and will hereafter be referred to as faults. Active reverse faults are shown in blue, while active normal faults are represented in red. The topographic slope ($\alpha$) is indicated by a colored line, located above the tectonic structure. Finally, since our simulations do not include creeping, the results are time-independent; i.e. 2 Myr of convergence at 5 cm. yr$^{-1}$ yields the same deformation as 1 Myr at 10 cm.yr$^{-1}$. For reasons of general applicability, we chose to describe the progress of the simulation in terms of amount of shortening (displayed in the inset at the bottom right of each frame) rather than time.

Just before the subduction of the first seamount ($\delta = 42.1$ km), the wedge is composed of a classic sequence of thrust at compressional critical state (Davis et al., 1983). Landward of the seamount, a small
<figure><img src="image_5.png" /><figcaption>**Fig. 2.** Subduction of the two first seamounts for reference simulation (M1). Seamounts subduction trigger gravitational collapse, basal erosion, slices of undeformed sediments both landward ($S_{lw}$) and seaward ($S_{sw}$), and the formation of a basin filled with syn-tectonic sediments. Inherited structures from the previous seamount are deformed.</figcaption></figure>

undeformed sediment slice ($S_{lw}$) is visible. The slice is composed of sediments located between the seamount and the last thrust (i.e., the backthrust-forethrust root). Unaffected by deformation and therefore softening, this slice exhibits greater resistance compared to the surrounding wedge material.

Once the seamount starts to subduct ($\delta = 58.3$ km), the weak décollement layer is suddenly interrupted by a higher frictional zone (30° instead of 5°). As a result, the décollement is abandoned and deformation propagates upward through the strata. This newly formed thrust located above the seamount acts as the new basal fault, which induces an increase of the basal friction, from 5° (décollement friction) to 10° (friction of deformed sediments), and an increase of the topographic slope. This new basal fault also leads to a steepening of the

thrust sequence landward of the seamount.

To monitor the evolution of the length of the area disturbed by the subduction of the seamount, we define $L_{dist}$ as the region where the topographic slope exceeds the internal friction angle of the prism, i.e., 10° (for deformed sediments) (Fig. 2).

$L_{dist}$ increases from 9.3 km to 12.2 km when the shortening increases from $\delta = 58.3$ km to $\delta = 60.4$ km. As $L_{dist}$ grows, the disturbed portion of the prism becomes unstable and undergoes a gravitational collapse, evidenced by the detachment fault at its base (active normal fault shown in red, when $\delta = 60.4$ km). Despite their collapse, the stacked sediments retain a degree of structural coherence. The sediments remobilized during the gravitational collapse are distinguished by a different (greenish) color. They are characterized by a seaward horizontal ve-
locity greater than 3 cm·yr⁻¹. While the deformation front remains inactive, sediments at the front of the prism are underthrusted without undergoing deformation.

During burial, as observed after $\delta = 69.9$ km, the subducting seamount continues to cause an increase in topographic slope ($L_{dist}$ reaches 16.5 km), triggering gravitational collapse as it progresses. At the same time, the deformation front is no longer disturbed and accretion resumes at the front of the wedge. The sediments underthrusted below the basal fault atop the subducting seamount, form a slice located seaward of the seamount ($S_{sw}$, as shown after $\delta = 69.9$ km). Like $S_{lw}$, $S_{sw}$ remains unaffected by softening and is therefore more resistant than the surrounding wedge material. As indicated by the deformation rate after $\delta = 69.9$ km, the reactivation of the deformation front also involves the re-engagement of the basal décollement seaward of the seamount. As a result, the $S_{sw}$ slice becomes decoupled from the model base and is incorporated to the upper plate. In contrast, the seamount remains coupled to the base. Consequently, following the reactivation of the deformation front, the seamount moves faster than $S_{sw}$, leading to extension between these two geological features, as evidenced by the normal fault (in red) separating the slice from the seamount. This process induces a basal erosion of the accretionary prism (Lallemand et al., 1994), as $S_{lw}$ is removed from the upper plate and transported down dip with the oceanic plate. Finally, after the reactivation of the deformation front, extension is observed at both surface and depth.

The black dashed line (Fig. 2a) indicates the location of the deformation front prior to the first seamount arrival. We can notice that, although 17.8 km of shortening have been accommodated since the 1st seamount arrival, the length of the prism has decreased due to the combined effects of the topographic slope increase caused by the subducting seamount and the interruption of the frontal accretion of sediments.

Before the second seamount arrival ($\delta = 136.9$ km), we can observe the structural features left by the subduction of the first seamount: a large mass transport deposit, a sedimentary basin at its front, and an undeformed slice at its base. The basin, filled with syn-tectonic sediments, results from the migration of the deformation from the gravitational collapse region to the deformation front ($\delta = 69.9$ km). The syn-tectonic sediments are represented in various shades of green and yellow, indicating their deposition age (here, the amount of shortening). Thus, sediments deposited at the bottom of the basin are coeval with the reactivation of the deformation front.

The first seamount is disintegrated by the backstop and progressively evacuated from the model through the lower right corner, where a non-zero velocity is imposed (see methodology section).

The second seamount induces similar deformation in the prism ($\delta = 136.9$ km, 150.7 km, 155.8 km), involving chronologically the formation of a slice landward from the seamount ($S_{lw}$), an increase of the topographic slope and a gravitational collapse during the seamount burial, a basal erosion of the prism, the formation of a slice seaward of the seamount ($S_{sw}$) made of underthrusted and undeformed sediments.

The second seamount deforms the inherited structures left by the previous seamount along its path: the mass transport deposits and the basin slice are compressed, while the $S_{sw}$ slice overthrusts the seamount. Despite these deformations, all the structures remain identifiable.

At the front of the wedge, accretion resumes ($\delta = 155.8$ km). Frontal compression is primarily accommodated seaward of the 2nd $S_{sw}$ slice. The slice, composed of more resistant sediments than those of the surrounding wedge, acts as a dynamic backstop against which new sediments are accreted ($\delta = 176.4$ km). As a result, the prism exhibits a three-part structure (from seaward to landward): 1) a sequence of thrust, 2) the inherited structures of the second seamount, consisting of a slice and a basin, and 3) a chaotic zone with mass transport deposits at the surface and the deformed inherited structures from the subduction of the 1st seamount.

We can also notice that due to the normal fault located on the seaward flank of the seamount, the $S_{sw}$ slice and the seamount get

separated by 17 km after $\delta = 176.4$ km.

## 3.2. – smaller seamounts

Simulation M2 is similar to simulation M1 but features smaller seamounts (Table 2). Like Fig. 2, Fig. 3 illustrates the subduction of the first two seamounts in simulation M2, divided into six stages.

Before the first seamount arrival ($\delta = 42.8$ km), as in M1, the prism forms a typical accretionary wedge and exhibits a slice ($S_{lw}$) landward of the seamount, composed of undeformed and more resistant sediments.

Once the seamount starts to subduct ($\delta = 53.8$ km), the prism continues to display a structure similar to M1. The seamount disrupts the front of the prism, forcing the basal fault to propagate upward through the strata and increasing the topographic slope. Accretion ceases, and sediments are underthrusted. The disturbed area ($L_{dist}$) extends over 7.87 km. We also observe a steepening of the thrust sequence formed before the seamount arrival on the landward side of the seamount.

After only 17.1 km of subduction, the seamount gets decapitated ($\delta = 57.1$ km). The base of the seamount keeps being coupled to the base of the model, and is dragged down by the oceanic crust, as illustrated in simulation M2_withcrust in Supplementary Material, while its upper portion is trapped within the prism. Once decapitated, the basal décollement, as well as the deformation front, are reactivated. Unlike M1, no gravitational collapse is observed. The premature reactivation of the deformation front, combined to the reduced size of the seamount, prevented $L_{dist}$ from reaching a sufficient length for the prism to become unstable, thereby limiting the size of the slice $S_{sw}$ (13.6 km for M1 vs 9.7 km for M2). Consequently, the retreat of the deformation front consecutive to the first seamount subduction is less pronounced in M2 than M1.

Before the subduction of the second seamount ($\delta = 13.8$ km) as for M1, structures inherited from the first seamount subduction are observable within the prism: a slice of underthrusted sediments covered by a basin filled with syn-tectonic sediments. Due to the absence of gravitational collapse, no mass transport deposits are observed at the surface.

Like the first seamount, the second seamount is associated with landward ($S_{sw}$) and seaward ($S_{lw}$) slices, and is decapitated during burial, without triggering gravitational collapse ($\delta = 138.6$ km).

Since both the slice $S_{sw}$ and the upper part of the seamount are decoupled from the base of the model when the deformation front is reactivated, no extension occurs between the upper part of the seamount and $S_{sw}$, explaining why both features remain connected after $\delta = 150.1$ km. Since the seamount no longer induces deformation, the inherited structures from the first seamount are well preserved, in contrast to M1. The overall structure of the wedge is finally composed of a sequence of thrusts, interspersed with basins and slices inherited from past seamount that have been decapitated.

## 3.3. – relationships between uplift, gravitational collapse, basal erosion and seamount decapitation

Fig. 4 illustrates the length of the disturbed area ($L_{dist}$) (length along which the topographic slope exceeds 10°, the lower friction of deformed sediments) caused by the second seamount subduction in simulations M1 and M2 as a function of the amount of shortening. To make our results more universally comparable, $L_{dist}$ is normalized by the width of the seamount ($w_{sm}$) and $\delta$ is adjusted to be zero at the onset of seamount burial.

The stars correspond to the values at the time steps shown in Fig. 2b and Fig. 3b. The dots highlight key moments in the deformation of the wedge.

For M1, we observe that $L_{dist}$ increases with seamount burial. The gravitational collapse begins once $L_{dist}$ exceeds $3.1 \times w_{sm}$ (Lcol). $L_{dist}$ and the length of the gravitational collapse zone continue to increase up to around $4.1 \times w_{sm}$. Once deformation resumes at the front, $L_{dist}$ remains stable and the domain undergoing gravitational collapse propagates
<figure><img src="image_6.png" /><figcaption>**Fig. 3.** Subduction of the two first seamounts for simulation M2. Smaller seamounts are decapitated. The seamount burial does not induce any collapse, but do form a basin and slices both landward ($S_{lw}$) and seaward ($S_{sw}$). Inherited structures from previous seamounts are preserved.</figcaption></figure>

through the wedge. Mechanically, the deformation front is reactivated when it becomes easier to cut downward within the sediment pile at the front of the prism rather than maintaining basal thrust activity within the strata. This scenario is referred as **basal erosion** and we define $L_{be}$ as the value of $L_{dist}$ at which the deformation front is reactivated which occurs at $4.1 \times w_{sm}$ and is reached after 27 km of shortening.

For M2, the increase in $L_{dist}$ is prematurely interrupted at $2.2 \times w_{sm}$ ($L_{dec}$) and 10.1 km of shortening when the seamount is decapitated. Mechanically, a seamount is decapitated when it becomes easier to reactivate the décollement layer despite the frictional heterogeneity induced by the seamount, rather than maintaining basal thrust activity within the strata. This scenario, allowing reactivation of the deformation front, is referred as **decapitation**. Since the decapitation occurs before $L_{dist}$ reaches the minimal length for collapse, no gravitational collapse is associated to the subduction of this seamount.

To investigate the effect of the décollement friction, we performed additional simulations: sF1, IF1 similar to M1, and sF2 and IF2 simulations, similar to M2, where we varied the friction of the décollement (2° and 8° respectively, Table 2). As described in previous studies (Davis et al., 1983; Huiqi et al., 1992; Ruh et al., 2012), the lower the friction,

the smaller the topographic angle and the longer and narrower the wedge (see video in supplementary material).

In sF1 and IF1, deformations are similar to those observed in M1, namely: 1) seamounts are not decapitated 2) gravitational collapse and basal erosion are observed and 3) inherited structures from previous seamounts are deformed. For both simulations, no trend is observed between the décollement friction and $L_{col}$. In all simulations, the disturbed part of the wedge slides over a newly formed basal thrust on top of the seamount. Hence, the deformation of this part of the wedge depends on the friction of this new thrust, corresponding to weakened sediments, explaining the lack of correlation with the décollement friction.

$L_{be}$ increases with the décollement friction (Fig. 5). The greater the décollement friction, the more difficult it becomes to cut downward within the sediment pile, explaining why $L_{be}$ increases with décollement friction.

For sF2 and IF2, deformations are similar to M2: seamounts are decapitated and collapse is not observed (see video in supplementary material), except for the first seamount in IF2 where $S_{lw}$ is inexistent (see next paragraph for explanations). $L_{dec}$ increases with décollement friction.
<figure><img src="image_7.png" /><figcaption>**Fig. 4.** Length of the disturbed area ($L_{dist}$) caused by the second seamount subduction in simulations M1 and M2 as a function of the amount of shortening. $L_{dist}$ is normalized by the width of the seamount ($w_{sm}$). $\delta$ is adjusted to be zero at the onset of seamount burial. No collapse is observed if the **decapitation** occurs before $L_{dist}$ reaches $L_{col}$. In contrast, gravitational collapse always occurs in the basal erosion scenario ($L_{dist} > L_{col}$).</figcaption></figure>

<figure><img src="image_8.png" /><figcaption>**Fig. 5.** $L_{be}$ and $L_{col}$ as a function of a) décollement friction b) elastic thickness of the lithosphere (EET) c) $L_{dec}$ as a function of décollement friction.</figcaption></figure>

tion (Fig. 5c), as higher friction hinders the reactivation of the basal thrust at the base of the model.

We also compared simulations with different lithospheric elastic thickness (EET), sE1 and lE1 similar to M1 but with EET = 10 km and 40 km, respectively (Table 2). As described in Ruh (2020), a higher EET leads to a lower basal slope, resulting in a smaller topographic angle and a longer, narrower wedge (see video in the supplementary material). Deformations in sE1 and lE1, like for sF1 and lF1, are similar to those observed in M1. According to Fig. 5b, $L_{be}$ increases with EET but no trend is observed between $L_{col}$ and EET.

To illustrate the importance of the slice $S_{lw}$, we performed a new simulation M2' adjusting the initial position of the first seamount to reduce $S_{lw}$ to zero (Fig. 6, $\delta = 50.3$ km shortening). In this case, the seamount is not decapitated (Fig. 6, $\delta = 58.1$ km) but induces a **basal erosion** of the prism., similar to M1, with an active décollement both seaward and landward of the seamount, and extension occurring between the slice and the seamount. The same scenario is observed in simulation lF2 (identical to M1 but with 8° décollement friction; see Table 2 and video in the Supplementary Material).

As a consequence, we here show that seamounts can be decapitated. Their decapitation depends not only on their size, but also on the presence of an undeformed and resistant slice at their front. The development of this slice depends on the timing of the seamount's arrival at the deformation front relative to the backthrust-forethrust succession ($S_{lw}$ length).

Characteristics length ($L_{be}$, $L_{dec}$, and $L_{col}$) can only be measured when they are reached. However, theoretically, we can assume that for each seamount, depending on the geological parameters and $S_{lw}$ size, values for $L_{be}$, $L_{dec}$, and $L_{col}$ are associated.

When the wedge is disturbed, there is a competition between **basal erosion** and **decapitation** scenarios for the reactivation of the deformation front. If $L_{dec} < L_{be}$, the seamount will be decapitated; otherwise, the deformation front will be reactivated after a basal erosion of the prism. In the simulations we performed, only smaller seamounts are decapitated. This can be explained by the fact that larger seamounts introduce more significant frictional heterogeneity, which increases the average friction in the disturbed part of the wedge. This leads to an increase in $L_{dec}$, resulting in $L_{dec} > L_{be}$.

Similarly, $L_{col} < \min(L_{dec}, L_{be})$, defines conditions for collapse to occurs. In our study, only two cases are observed: **basal erosion** with collapse ($L_{col} < L_{be} < L_{dec}$) or **decapitation** without collapse ($L_{dec} < \min(L_{thr}, L_{col})$). Although the explored parameter range did not capture all possible scenarios, two additional outcomes are theoretically conceivable for $L_{col} < L_{dec} < L_{be}$ **basal erosion** could occur without collapse, while $L_{dec} < \min(L_{be}, L_{col})$ should result in **decapitation** with collapse.

### 3.4. – stress state

Fig. 7 shows the tectonic overpressure, defined as the difference
<figure><img src="image_9.png" /><figcaption>**Fig. 6.** Subduction of the first seamount for simulation M2'. M2' is the same simulation as M2, but with an adjusted initial position of the first seamount to reduce $S_{lw}$ to zero. This results in the small seamount not being decapitated.</figcaption></figure>

between the actual and lithostatic pressures in the prisms, as well as the orientation of the largest principal stresses ($\sigma_{max}$, black lines) for M1 (Fig. 7a) and M2 (Fig. 7b) simulations. The actual pressure is the mean of principal stresses.

The three stages of simulation M1 presented (Fig. 7b) correspond to stages 1, 4 and 7 of Fig. 2. Before the arrival of the first seamount ($\delta = 42.1$ km), the basal layer of the prism is moderately overpressurized (around 15 MPa). The high overpressure (around 55 MPa) found at the base of the backstop is driven by the boundary conditions imposed on the right edge of our model (zero velocity), and thus not interpreted.

During the seamount subduction and the collapse stage ($\delta = 69.9$ km), a strong tectonic overpressure is found on the landward side of the seamount along the slice $S_{lw}$ (around 50 MPa), with horizontally oriented $\sigma_{max}$. On the seaward side of the seamount, the whole slice $S_{sw}$ is affected by a tectonic underpressure (around -30 MPa) and vertically oriented $\sigma_{max}$. The subduction of the second seamount induces a similar pattern in the stress state of the wedge: overpressure within slice $S_{lw}$ and underpressure within slice $S_{sw}$ (Fig. 7a, $\delta = 155.8$ km). The stresses associated with the $S_{sw}$ slice inherited from the first seamount subduction, display a particular pattern (Fig. 7a, $\delta = 155.8$ km), with tectonic underpressure along the décollement and the basal part of the slice, and a strong tectonic overpressure at the top of the slice and at the base of the syn-tectonic basin formed above.

For the smaller seamounts (M2, Fig. 7b), before being decapitated

(Fig. 7b, $\delta = 55.5$ km), the seamount induces a slightly higher tectonic overpressures on its landward side (around 30 MPa) and slightly lower underpressure seaward (around -15 MPa).

Once decapitated (Fig. 7b, $\delta = 61.5$ km), the head of the seamount is decoupled from the base of the model and no longer influences the stress state of the wedge. However, the base of the seamount continues subducting and acting as a frictional discontinuity. As shown in the zoomed-in view, tectonic overpressure is observed landward of this discontinuity and underpressure seaward, which results in an underpressure and vertical $\sigma_{max}$ within $S_{lw}$. The same scenario applies for the second seamount (Fig. 7b, $\delta = 138.0$ km, 143.7 km).

# 4. – discussion

## 4.1. – *structures related to seamount subduction*

Except for well-instrumented margins, links between shallow SSEs and seamounts remain difficult to establish, partly due to challenges in identifying subducted seamounts (Bassett and Watts, 2015). We therefore propose structural and morphological criteria to locate both active and past subducted seamounts.

The pioneer work of Dominguez et al. (1998, 2000) was the first to describe the characteristic features of seamount subduction: a trench re-entrant, uplift of the wedge followed by its subsidence with
<figure><img src="image_10.png" /><figcaption>**Fig. 7.** Tectonic overpressure in the prism (defined as the difference between the actual and lithostatic pressures) and orientation of the largest principal stresses ($\sigma_{max}$) for a) M1 and b) M2 simulations.</figcaption></figure>

Before the reactivation of the deformation front, tectonic underpressure and vertically oriented $\sigma_{max}$ are observed within slice $S_{sw}$, while tectonic overpressure and horizontally $\sigma_{max}$ are present within slice $S_{lw}$. Immediately after reactivation, tectonic underpressure develops within $S_{lw}$ in the M2 simulation (**decapitation** scenario), whereas in the M1 simulation (**basal erosion** scenario), previous conditions persist.

gravitational collapse, and the underthrusting of undeformed sediments accumulated seaward of the seamount. These features were further confirmed by both analogue and numerical modelling (Dominguez et al., 1998, 2000; Noda et al., 2020; Ruh, 2016; Ruh et al., 2016; Morgan et al., 2017; Wang et al., 2021; Okuma et al., 2022; Miyakawa et al., 2022; Pajang et al., 2022a). The same features are obtained in the case of a buried seamount, provided that the decollement within the sedimentary pack lies deeper than seamount's summit (Pajang et al., 2022a).

A seamount acts as a frictional discontinuity along the décollement, forcing the basal fault to migrate upward within the strata, thereby increasing the friction and the slope of the basal fault (Wang et al., 2021). By disturbing the accretionary wedge, a seamount can trigger gravitational collapse (Fig. 2). Gravitational collapse was previously thought to occur only when seamounts protrude above the sediment cover (Ruh, 2016). However, we here show that even initially buried seamounts can induce such collapse. We also demonstrate that collapse actually occurs when a certain length of the wedge is disturbed ($L_{col}$). In our simulations, when **decapitation** scenario causes reactivation of the deformation front, seamount is sheared before reaching this minimum length ($L_{col}$) and therefore do not induce any collapse.

A slice of underthrusted sediments forms up-dip for both small and large seamounts ($S_{sw}$), as already discussed in previous studies (Dominguez et al., 1998, 2000; Wang et al., 2021; Miyakawa et al., 2022; Pajang et al., 2022ab). A slice of underthrusted sediments also forms downdip of the seamount ($S_{lw}$) as also described by Wang et al. (2021) with analogue modeling. We here specify that its extent depends on the timing of the arrival of the seamount relative to the timing of the frontal thrust activation (Mary et al., 2013a, 2013b). Depending on $S_{lw}$ size, small seamounts may or may not be decapitated.

By taking surface processes (erosion and sedimentation) into account, we highlight the formation of basins filled with syn-tectonic sediments, associated with the subduction of seamounts above slice $S_{sw}$.

## 4.2. – decapitation/exhumation?

Based on scarce evidence of exhumed seamounts, the decapitation of seamount during their subduction has been systematically disregarded, in favor of a *break through the upper plate* hypothesis (Wang and Bilek, 2014). However, some recent studies have suggested that exhumed seamounts might have been significantly underestimated (Bonnet et al., 2024). In our simulations, **small seamounts can be easily sheared** and are then evacuated from the model through the lower right corner. As a result, our simulations indicate that the remnants of seamount decapitation are not exhumed but subducted deeper. However, this behavior is strongly influenced by the imposed boundary conditions. Consequently, addressing the question of exhumation with our "sandbox-like" setup does not seem relevant. Nevertheless, larger scale configurations, such as Menant et al. (2020) setup, would be more suitable for investigating exhumation.

## 4.3. – comparison with natural examples

To demonstrate that our simulations correctly capture structures associated to seamount subduction, we now compare our results with two seismic profiles located along Nankai (KIN028, Flores et al. (2025); Nakamura et al. (2022), Fig. 8a) and Hikurangi (05CM-04, Barker et al. (2018), Fig. 8c). According to Flores et al. (2025), Fig. 8a shows one clearly defined seamount and a deeper bathymetric high. From North to South, we observe the following:

* Landward of the bathymetric high: a thrust sequence above a suspected underplated unit. The deformation of the thrust sequence likely results from the compression imposed by the subducting bathymetric high as observed in simulation M1 at $\delta = 69.9$ km or M2 at $\delta = 134.8$ km (Fig. 2 and 3). The underplated unit could correspond to the $S_{sw}$ slice of a former relief
<figure><img src="image_11.png" /><figcaption>Fig. 8. a) Interpretation of a seismic transect along the Nankai margin (KIN028) from Flores et al. (2025), illustrating the successive subduction of a topographic high followed by a seamount, along with the associated deformation.</figcaption></figure>

b) Seismic transect (05CM-04 from Barker et al. (2018)) overlaid on P-wave velocity model of the Papaku seamount along the Hikurangi margin (New-Zeland) compared to c) simulation M1 at $\delta = 60.4$ km. Inactive faults are identified using accumulated strain ($\epsilon_p > 2$).

The description of the highly reflective zone (HRZ), interpreted as a zone of high pore-fluid pressure hosting SSEs, is consistent with our $S_{lw}$ description.

* Seaward of the bathymetric high: a deformed slice starting at a trench slope break and located atop a low topographic slope. This feature corresponds to the $S_{sw}$ slice of the bathymetric high, as observed in simulation M2 at $\delta = 134.8$ km (Fig. 3). The proximity of

the $S_{sw}$ slice to the bathymetric high may reveal that the original relief has been decapitated.

* On top of the seamount: a deformed thrust sequence with some mass transport deposit. Both the $S_{sw}$ slice and the thrust sequence have most likely been deformed by the subducting seamount, similarly to
the second seamount in simulation M1 between $\delta = 155.8$ km (Fig. 2), although in this case, the gravitational collapse appears to be limited.

* Seaward of the seamount: a slice of underthrust sediments, which corresponds to the $S_{sw}$ slice of the seamount. This is followed by the thrust sequence formed by the reactivation of the deformation front.

Fig. 8b compares an annotated seismic line of the Papaku seamount from Barker et al. (2018) with the M1 simulation at the early stage of the first seamount subduction ($\delta = 60.4$ km). The seismic line shows that when a seamount is subducted, the basal thrust propagates upwards through the strata. On both sides of the seamount, we recognize the $S_{sw}$ and $S_{lw}$ slices of underthrust and underformed sediments. Unlike Barker et al. (2018), the thrusts faults observed above the detachment are inactive in our simulation and no mass transport deposit can clearly be identified. This discrepancy may be attributed to geometrical and material heterogeneities not taken into account in our model. Additionally, part of the sediment we identify as a mass transport deposit in the simulations may not be recognized as such in natural settings, as it retains a degree of structural coherence.

### 4.4. – comparison of the long-term stress state with slip behavior

High pore fluid pressure reduces the effective normal stress on faults, thereby increasing the nucleation size and favoring the occurrence of slow slip events (SSEs) (Ampuero and Rubin, 2008; Liu and Rice, 2007, 2009). When shallow SSEs are relocated onto seismic profiles, they systematically coincide with regions showing high vp/vs ratios or zones of strong seismic reflectivity (Wallace et al., 2009; Bell et al., 2010; Wallace and Beavan, 2010; Saffer and Tobin, 2011; Saffer and Wallace, 2015), both indicative of fluid-rich, overpressured environments. Although seismic data, particularly from active surveys, are often acquired at different times than the SSE recordings, this spatial correlation persists. It therefore suggests that the elevated pore fluid pressures imaged geophysically are not transient fluctuations associated with individual shallow SSEs, but rather represent a long-lived, quasi-permanent feature.

During subduction, fluids are initially released through compaction of under consolidated sediments, followed by clay dehydration reactions (e.g., Saffer and Tobin, 2011). However, the development of near lithostatic pore fluid pressure requires that sediments remain undrained.

Under extensional conditions, however, such high pore pressures can only be sustained when the differential stress remains very small, below the cohesive strength of rocks. Once tensile failure occurs, vertical fractures open (Sibson, 1981, 1990) and act as drains, allowing fluids to escape periodically and thus preventing the maintenance of long-term overpressure. In contrast, under compressional stress, tensile failure generates subhorizontal mode-I cracks that inhibit vertical drainage (Sibson, 1981), leading to the progressive lateral expansion of over-pressured regions.

Moreover, even in a compressive regime, fault reactivation (which enhances fluid drainage) typically occurs before lithostatic pore fluid pressure are reached (Sibson, 1990, 2017). Therefore, undeformed sediments in a compressional tectonic setting provides favorable conditions for the buildup of long-term lithostatic pore fluid pressure.

As shown by Ruh et al. (2016), and our study, tectonic overpressure builds on the downdip side of the seamount and underpressure develops up dip. Additionally, along the $S_{sw}$ frontal slice, $\sigma_{max}$ is vertically oriented, indicating extensional conditions while along $S_{lw}$, $\sigma_{max}$ is vertical (compressional conditions).

Therefore, $S_{lw}$ slices, composed of underthrusted, undeformed and underconsolidated sediments, and subjected to a compressional tectonic overpressure are a favorable environment to build long-term high pore fluid pressures.

In Nankai, very low frequency earthquakes and tremors have been reported on the landward side of seamounts, along underthrusted units

of fluid-rich sediments compressed by a newly subducting seamount (Miyakawa et al., 2022; Flores et al., 2025). For the seamount depicted in Fig. 8a, a low pore fluid pressure has been estimated at its front based on P-wave velocity and taper angles (Flores et al., 2024), whereas a low-velocity zone indicating a high pore fluid pressure has been imaged on its landward side (Tonegawa et al., 2017; Flores et al., 2025). Along Hikurangi, the largest slip magnitude of SSE occurs downdip of the Papaku seamount, along a highly reflective zone (HRZ), interpreted as 3.0-km-thick fluid-rich subducting sediments (Bell et al., 2010, 2014; Barker et al., 2018; Bangs et al., 2023) (Fig. 8b). This HRZ zone, interpreted as a zone of high pore fluid pressure, coincides with our suggested $S_{lw}$ slice and supports our hypothesis. Therefore, $S_{lw}$ slices could provide favorable conditions for the buildup of near lithostatic pore fluid pressure and slow slip events nucleation.

Along Hikurangi, the seamount itself corresponds to a sharp decrease in slip and to the locus of tectonic tremors (Barker et al., 2018). A model explicitly resolving pore fluid pressure and slip behavior would be necessary for a better understanding of the occurrence and propagation of these events.

### 5. – conclusion

The subduction of buried seamounts deforms the accretionary prism. Undeformed sediments are preserved both landward ($S_{lw}$) and seaward ($S_{sw}$) of the seamount. When surface processes are included, basins filled with syn-tectonic sediments also form.

We identify two distinct mechanisms for deformation front reactivation, each with specific structural consequences:

* **Basal erosion:** The seamount remains undeformed. If the seamount-disturbed zone ($L_{dist}$) is sufficiently wide, gravitational collapse occurs. When multiple seamounts subduct sequentially, only the most recent structures remain distinguishable; older ones are progressively overprinted and deformed.

* **Decapitation:** Smaller seamounts can be decapitated. In this case, no gravitational collapse occurs, and structural inheritances—$S_{sw}$ slices and syn-tectonic basins—are well preserved.

We demonstrate that our structural results can be used to interpret seismic lines by comparing them with seismic profiles from Nankai and Hikurangi.

Prior to deformation front reactivation, tectonic overpressure builds up landward of the seamount, while underpressure develops seaward. This results in vertically oriented $\sigma_{max}$ within $S_{sw}$ and horizontal $\sigma_{max}$ within $S_{lw}$.

We then discuss the impact of the tectonic structures and stress state in the wedge on the long-term pore fluid pressure, without explicitly modeling fluid flow. Although $S_{sw}$ and $S_{lw}$ are not fault-drained, their distinct stress fields have potentially contrasting effects: in $S_{sw}$, vertical $\sigma_{max}$ promotes fluid drainage; in $S_{lw}$, horizontal $\sigma_{max}$ favors long-term high pore fluid pressure development.

Our study provides mechanically based explanation for the occurrence of long-lived high pore fluid pressure zones landward of seamounts, where shallow SSEs are observed to nucleate, as along the Hikurangi margin.

#### Funding

This study was funded by the ANR MechaThrust.

#### CRediT authorship contribution statement

**Alexis Gauthier:** Writing – original draft, Software, Methodology, Investigation, Conceptualization. **Nadaya Cubas:** Writing – review & editing, Visualization, Validation, Supervision, Funding acquisition, Conceptualization. **Laetitia Le Pourhiet:** Writing – review & editing,
Validation, Supervision, Software, Methodology.

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.

Supplementary materials

Supplementary material associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2025.119767.

Data availability

Data will be made available on request.

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