# Fluids, faulting and earthquakes in the brittle crust: recent advances and new challenges

Olivier Fabbri, Hugues Raimbourg, Henri Leclère

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Olivier Fabbri, Hugues Raimbourg, Henri Leclère. Fluids, faulting and earthquakes in the brittle crust: recent advances and new challenges. Comptes Rendus. Géoscience, 2024, 356 (S2), pp.423-466. 10.5802/crgeos.259. hal-04758533

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# Comptes Rendus
## Géoscience
### Sciences de la Planète

Olivier Fabbri, Hugues Raimbourg and Henri Leclère

**Fluids, faulting and earthquakes in the brittle crust: recent advances and new challenges**

Volume 356, Special Issue S2 (2024), p. 423-466

Online since: 21 May 2024

**Part of Special Issue:** Geodynamics of Continents and Oceans – A tribute to Jean Aubouin

**Guest editors:** Olivier Fabbri (Université de Franche-Comté, UMR CNRS 6249, Besançon), Michel Faure (Université d'Orléans-BRGM, UMR CNRS 7325, Institut des Sciences de la Terre, Orléans), Jacky Ferrière (Université de Lille, faculté des Sciences), Laurent Jolivet (Sorbonne Université, ISTE, UMR 7193, Paris) and Sylvie Leroy (Sorbonne Université, CNRS-INSU, ISTE, Paris)

https://doi.org/10.5802/crgeos.259

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This article is licensed under the
CREATIVE COMMONS ATTRIBUTION 4.0 INTERNATIONAL LICENSE.

http://creativecommons.org/licenses/by/4.0/
Research article

Geodynamics of Continents and Oceans – A tribute to Jean Aubouin

# Fluids, faulting and earthquakes in the brittle crust: recent advances and new challenges

Olivier Fabbri<sup>●,*,a</sup>, Hugues Raimbourg<sup>●,b</sup> and Henri Leclère<sup>●,a</sup>

<sup>a</sup> UMR 6249, université de Franche-Comté, 16 route de Gray, 25030 Besançon
cedex, France

<sup>b</sup> UMR 7327, universite d'Orléans, 1A rue de la Férollerie, 45071 Orléans
cedex 2, France

E-mail: olivier.fabbri@univ-fcomte.fr (O. Fabbri)

**Abstract.** Interactions between fluids and deformation are widespread in the brittle crust. As experimentally shown, a high pore fluid pressure $p_f$ can fracture intact rocks or reactivate pre-existing fractures. The preference of reactivation over the formation of a new fracture depends on the orientation of the pre-existing fracture with respect to the stress axes and on $p_f$. In nature, the predominant reactivation of misoriented pre-existing faults rather than the formation of new faults with more favorable orientations suggests that pressurized fluids are present in the brittle crust. There is a large body of evidence indicating that supra-hydrostatic $p_f$ contributes to the reactivation of low-angle thrust faults or normal faults. Conversely, supra-hydrostatic $p_f$ values are less common along vertical or steeply dipping plate boundary transform faults or intra-continental strike-slip faults. If these faults are severely misoriented with respect to the ambient stress field, their reactivation may not be due to supra-hydrostatic $p_f$ but to other mechanisms such as shear-enhanced compaction or thermal pressurization. Supra-hydrostatic $p_f$ also plays a role in the nucleation or propagation of seismic ruptures in the continental or oceanic crust, and in subducting slabs in convergent margins, as reported for aftershocks, swarms, slow earthquakes, and to a lesser extent for major earthquakes. Lastly, increase or decrease of $p_f$ in depth due to human activities such as hydrocarbon extraction, dam impoundment, gas storage or geothermal energy production result in many cases in the inception or enhancement of seismic activity, adding clues in favor of a relationship between fluids and earthquakes.

**Keywords.** Fluid, Fault, Pore fluid pressure, Earthquake, Slow earthquake, Seismic swarm, Crack-seal vein.

Manuscript received 8 December 2023, revised 12 April 2024, accepted 15 April 2024.

## 1. Introduction

Whatever their nature, chemical composition or physical state, fluids play key roles in the evolution of the Earth's lithosphere, particularly regarding sediment diagenesis, volcanic activity, hydrothermal alteration, metamorphism, ore deposit formation,

among many other phenomena. Fluids are also important with regard to deformation, especially brittle deformation. In the upper brittle crust, fluids can circulate in deforming zones, namely fracture zones or fault zones, and can chemically or mechanically interact with host rocks. Such interactions are also conceivable in subduction zones, where the upper part of the subducting plate keeps a brittle behavior down to hundreds of kms of depth.

<footnote>* Corresponding author</footnote>
The role that fluids exert on rock deformation is of paramount importance for human societies. Indeed, fluid extraction (e.g., hydrocarbon industry) or fluid injection (geothermal energy, waste water disposal, gas storage) trigger or enhance seismic activities in otherwise aseismic or moderately active regions.

This review aims at bridging the gap between the mechanical interactions between fluids and rocks on one hand and the structural analysis of deformed rocks from the scale of the hand sample to that of the deforming plate boundaries. For space reasons, the chemical effects of fluids, although of considerable importance, will not be addressed here.

## 2. The mechanical interaction between fluids and rocks

### 2.1. The development of abnormally high pore fluid pressures in the upper crust

Production of fluid in depth, for instance by magmatic activity or by metamorphic or metasomatic reactions, can result in an influx of fluids. If there is a permeability barrier that prevents or hinders fluid escape, then the pore fluid pressure $p_f$ will increase. This increase can also be generated by pore volume reduction in low-permeability rocks. For instance, elastic crack relaxation following an earthquake (seismic pumping, see below Section 2.6) or inelastic pore compaction controlled by the applied shear stress (shear-enhanced compaction, see below Section 2.7) are two important mechanisms responsible for pore volume decrease. Beside such earthquake-related mechanisms, pore compaction can also be achieved during diagenesis [Walder and Nur, 1984, Wang et al., 2022]. As attested by measurements in deep boreholes, supra-hydrostatic $p_f$ values actually develop below impermeable barriers or seals, and can reach the value of the lithostatic stress or even exceed it [Yerkes et al., 1985, Powley, 1990, Neuzil, 1995, Zencher et al., 2006].

### 2.2. Intact rock fracturing, stress corrosion, hydrofracturing

Classical rupture criteria for intact rocks in the upper crust are inspired by experimental rock mechanics and state that, as long as the stress tensor (whose components are $\sigma_1$, $\sigma_2$, and $\sigma_3$, with the convention

$\sigma_1 > \sigma_2 > \sigma_3$) acting on the rock does not satisfy any criterion, no fracture will form. To the contrary, once the stress tensor satisfies the criterion, microscopic fractures will propagate through the rock, eventually leading to the formation of a macroscopic fracture or fault.

Two rupture criteria are commonly used in the analysis of intact rock failure. The simplest is the linear Mohr-Coulomb criterion, which postulates that the propagation of macroscopic fractures happens when

$$ \tau = \mu_i \sigma_n + C $$

with $\tau$ being the tangential stress acting along the fracture, $\sigma_n$ the normal stress acting on the fracture, $\mu_i$ and $C$ being respectively the intrinsic coefficient of friction and the intrinsic cohesion, both depending on rock type. The more elaborated parabolic Griffith failure criterion can be expressed as

$$ \tau^2 + 4T\sigma_n = 4T^2 $$

where $T$ is the tensile strength of the intact rock, counted positive. Both Mohr-Coulomb and Griffith criteria can be combined in a composite criterion for which the parabolic criterion will apply to negative normal stress values and the linear criterion will apply for the positive normal stress values [Sibson and Scott, 1998; Figure 1].

In the absence of pore fluid in the rock, the convergence between the tensor and the criterion can be achieved only by increasing the differential (deviatoric) stress $\sigma_D = \sigma_1 - \sigma_3$, that is, either by decreasing $\sigma_3$, by increasing $\sigma_1$, or by acting on both $\sigma_3$ and $\sigma_1$. If a fluid is present inside the rock (so-called pore fluid), then two processes can lead to the formation of macroscopic fractures. In the first process, called stress corrosion, the fluid will chemically react at the atomic scale with the rock at the tip of a pre-existing loaded crack, so that the critical stress intensity factor at this location gets lower and lower. When the stress intensity factor at the tip of the crack reaches the fracture toughness of the altered rock, the fracture will propagate. This phenomenon is called sub-critical crack growth [Anderson and Grew, 1977, Atkinson, 1984, Brantut et al., 2013]. The second process requires that the pore fluid is elevated at a pressure $p_f$. If so, the components $\sigma_i$ of the tensor will then be changed into effective components $\sigma_i^* = \sigma_i - p_f$, and the rupture criterion is assessed using effective stress.
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**Figure 1.** The role of pore fluid pressure in fracturing intact rocks as represented in the Mohr space. The intact rock rupture criterion is composite, that is, it consists of the Griffith criterion for the negative normal stresses, and of the linear Mohr–Coulomb criterion for the positive normal stresses. $\sigma'_n$: effective normal stress acting on the surface ($\sigma'_n = \sigma_n - p_f$). (a) A pore fluid pressure $p_{f1}$ decreases the normal stresses and shifts the Mohr circle until it tangents the failure envelope. The orientation of the newly formed fracture with respect to the stress axes is optimal. (b) Due to a lower differential stress, a pore fluid pressure $p_{f2}$ displaces the Mohr circle until it tangents the failure envelope in the parabolic part, resulting in the formation of hybrid shear-extensional fractures.

**Figure 1.** (cont.) (c) Due to a still lower differential stress, a pore fluid pressure $p_{f3}$ displaces the Mohr circle until it tangents the failure envelope at $\tau = 0$, resulting in the formation of pure dilatant fractures.

If $p_f$ is increased enough, the effective stress tensor will satisfy the rupture criterion (Figures 1 and 2). Note that, unlike the $\sigma_i$ components of the tensor, $\sigma_D$ is not modified by any variation of $p_f$. A failure mode diagram [Cox, 2010] allows to estimate $\sigma_D$ and $p_f$ required for rock failure (Figure 3). In this diagram, $p_f$ is expressed by the pore fluid factor (also called pore fluid pressure ratio) $\lambda = p_f/\rho gz$, where $\rho gz$ is the lithostatic stress, $\rho$ the mean density of the rock column from the surface to depth $z$, and $g$ the gravity acceleration. The values of $\lambda$ are between 0.3–0.4 (hydrostatic $p_f$) and 1 (lithostatic $p_f$). In a Mohr–Coulomb diagram, if the tangent point between the Mohr circle and the envelope is located in positive $\sigma_n$ values, the rupture is said to be of shear failure type. This is illustrated in the failure mode diagram by the blue line on the failure envelope (Figures 1a and 3). If $\sigma_3^* = -T$, $T$ being the tensile strength of the intact rock, counted positive, the rupture is purely extensional (dilatant), and is illustrated in the failure mode diagram by the red line on the failure envelope (Figures 1c and 3). It is important to note that the use of extension here does not reflect an extensional tectonic regime (for which $\sigma_V = \sigma_1$, $\sigma_V$ being the principal vertical stress component, in an Andersonian framework). Intermediate locations correspond to hybrid extensional-shear failures illustrated in the failure mode diagram by the green line on the failure envelope (Figures 1b and 3).

Dilatant breccias form when the three effective principal stresses are negative ($\sigma_3^* < \sigma_2^* < \sigma_1^* < 0$), while pure dilatant breccias, that is, breccias without any preferred orientation of dilatancy (Figure 4), form when the stress tensor is hydrostatic, the common stress magnitude being the opposite of the tensile strength of the rock: $\sigma_1^* = \sigma_2^* = \sigma_3^* = -T$ [hydrostatic stress state; Cosgrove, 1995].

Mineralized extensional joints (tension gashes) are typical and widespread geological structures resulting from failure of intact rocks by pure extension (Figure 1c). Pure extension is achieved by the condition $\sigma_3^* = -T$, that is, $p_f = \sigma_3 + T$ [Hubbert and Willis, 1957, Secor, 1965, Hancock, 1985]. The
<figure><img src="image_4.png" /></figure>

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**Figure 2.** Competition between reactivation of pre-existing fractures and formation of a new fracture [after Handin, 1969, Sibson, 1985]. The failure envelope for intact rock follows a composite Griffith–Mohr–Coulomb criterion. The reactivation envelope is linear, and its cohesion is neglected. $\sigma'_n$: effective normal stress acting on the surface ($\sigma'_n = \sigma_n - p_f$). (a) The pre-existing fracture (black dot on the circle) is *favorably oriented* with respect to the $\sigma_1$ stress axis. The pore fluid pressure $p_{f1}$ is large enough to allow its reactivation. (b) The pre-existing fracture is *unfavorably oriented* with respect to the $\sigma_1$ stress axis, but the *misorientation is moderate*. The pore fluid pressure $p_{f2}$, larger than $p_{f1}$, is necessary to allow its reactivation. (c) The pre-existing fracture is *unfavorably oriented* with respect to the $\sigma_1$ stress axis, and the *misorientation is severe*.

**Figure 2. (cont.)** Due to an increasing pore fluid pressure, the Mohr circle shifts towards smaller normal stresses. It eventually tangents the intact rock failure envelope (grey dot) before the pre-existing fracture (black dot) can be reactivated. In this case, a new fracture (grey dot) with a favorable orientation is formed for a pore fluid pressure $p_{f3}$.

<figure><img src="image_7.png" /></figure>

**Figure 3.** Failure mode diagram for reverse and normal faulting based on the Mohr–Coulomb theory. Red, green and blue lines correspond respectively to pure extensional (dilatant) veins, hybrid veins and shear fractures. Oblique black lines correspond to the reactivation conditions for a favorably oriented fracture or a severely misoriented fracture. The diagram is drawn for a depth $z$ of 7 km, a rock density $\rho$ of 2.7, a friction coefficient $\mu$ of 0.75 and a tensile strength $T$ of 5 MPa. See text for explanation of points A, B and C.

regular and symmetrical layering of the minerals filling extension joints suggests a cyclical process with incremental opening stages alternating through time with fluid ingression and crystallization [Ramsay, 1980]. This so-called crack-seal mechanism, reminiscent of the cyclical nature of earthquakes, is often considered as a geological expression of the seismic cycle [see below and Raimbourg et al., 2021, 2022].
<figure><img src="image_8.png" /></figure>

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**Figure 4.** Purely dilatant breccias (hydraulic breccias) suggesting a hydrostatic stress state ($\sigma_1^* = \sigma_2^* = \sigma_3^* = -T$). (A) Saw-cut hand sample showing dark green serpentinite fragments cemented by white calcite along a paleo-detachment fault, Schistes Lustrés Zone, Queyras serpentinite, French Alps. Width of photograph: 25 cm. (B) River bed exposure of a dilatant breccia composed of low-porosity sandstone fragments (grey color) cemented by quartz (white color) along a vertical fault (white arrows), Oligocene Hyuga Group, Shimanto accretionary prism, Japan. (C) Saw-cut sandstone hand sample from nearby (B) exposure. Width of photograph: 15 cm.

Shear fractures result from a combination of extension and shear (hybrid failure). They classically form conjugate systems of *en échelon* fractures (Fig-

<figure><img src="image_11.png" /></figure>

**Figure 5.** Calcite-filled mode I and hybrid mineralized fractures associated with conjugate faults (arrows). The fractures are perpendicu- lar to the steeply dipping bedding surface. Cre- taceous limestones, Pyrenean foreland, Cor- bières, France (height of photograph: 2 m).

<figure><img src="image_12.png" /></figure>

**Figure 6.** Mode I quartz-filled fracture in low- porosity Oligocene sandstones, Shimanto ac- cretionary prism, Japan. The quartz fibers show an S shape, suggesting a component of dis- placement parallel to the fracture. Thickness of the vein: 15 mm.

ure 5), where the long axis of the mineral fibers indicates the combination of extension and shear (Figure 6).
## 2.3. Fracture reactivation vs. fracture neoformation

Unlike in laboratory experiments, rock masses in the brittle crust are fractured. Therefore, it is not the intact rock strength that will dictate the stability of rock masses, but rather the strength of discontinuity (fault or fracture) surfaces [Byerlee, 1978]. Intact rock rupture criteria are replaced by more realistic reactivation criteria. The simplest and most widely used criterion is the Mohr-Coulomb reactivation criterion [Handin, 1969], quite similar to the Mohr-Coulomb criterion for the intact rock, and which can be written as

$$ \tau = \mu\sigma_n + C. $$

The criterion postulates that the stability of fractured rock masses in the brittle crust is controlled by two basic parameters that can act in concert: (1) the coefficient of friction $\mu$ of the rock in contact or of the material filling the space between the two slipping surface walls, if any (e.g., fault gouge); (2) the normal stress component $\sigma_n$, acting perpendicularly to the discontinuity surface. The cohesion $C$ is usually assumed to be equal to zero, given its negligible value for most rocks. Handin [1969] showed experimentally that if a pre-existing fracture is unfavorably oriented with respect to the principal stress axes, a new fracture with a more favorable orientation will form, leaving the pre-existing one unreactivated. This can be illustrated by a construction in the Mohr diagram (Figure 2c) where the Mohr circle tangents the intact rock failure envelope before the point corresponding to the pre-existing fracture reaches the reactivation envelope.

## 2.4. The reactivation of moderately to severely misoriented fault surfaces

In the brittle crust, the experimentally demonstrated preference of fracture neoformation over preexisting misoriented fracture reactivation does not seems to be a general rule. Indeed, in many instances, moderately to severely misoriented fault surfaces are reactivated, while no propagation of new surfaces is observed. Several mechanisms are proposed to solve this paradox, among which the most frequently invoked are as follows.

(1) The fault zone consists of a continuous layer of weak rocks, for instance rocks whose rheology can

be regarded as plastic, visco-elastic, or elasto-plastic. This is the case with clay-rich or evaporite-rich strata (so-called décollement layers). Such a weak and continuous décollement layer will allow the displacement of allochton units over significant distances whatever the orientation of the fault zone with respect to the stress axes is favorable or not [Davis and Engelder, 1985, Weijermars et al., 1993, Costa and Vendeville, 2002, Vendeville et al., 2017].

(2) The fault surface is coated by low-friction ($\mu < 0.2$) material such as low-friction clay (e.g., montmorillonite) or talc [Morrow et al., 1992, 2017, Moore and Lockner, 2011, Boutareaud et al., 2012, Chen et al., 2017]. Leaving aside the stress corrosion process mentioned above, which concerns the nucleation of a mesoscopic fracture rather than reactivation of a macroscopic fracture or fault, fluids can weaken a fault by chemically reacting with the rocks or minerals along the damage zones or core zones of the fault. The chemically produced minerals can be frictionally weaker (so-called reaction softening). Examples of chemical weakening of fault rocks are described by White and Knipe [1978], Evans and Chester [1995], Wintsch et al. [1995], Gueydan et al. [2003], Collettini and Holdsworth [2004], Matsuda et al. [2004], Jefferies et al. [2006], Moore and Rymer [2007], Collettini et al. [2009b]. As indicated above, this reaction-softening effect will not be developed here.

(3) Along long-lived mature faults, damaged rocks inside the fault zone get their mechanical constants, such as their Young moduli or Poisson ratios, modified [Faulkner et al., 2006]. These modifications lead to a rotation of the stress axes in the vicinity of the fault [Healy, 2008]. The angle between $\sigma_1$ and the fault surface can thus be decreased and a very unfavorably oriented fault can become favorably oriented for reactivation. The studies of Provost and Houston [2001], Hardebeck and Michael [2004], Famin et al. [2014] suggest that vertical or steeply dipping strike-slip faults such as the San Andreas fault in California or the Median Tectonic Line in Japan can be reactivated following a rotation of the horizontal $\sigma_1$ axis to a more favorable value.

(4) The last mechanism, which is the main purpose of this paper, calls for the development of abnormally high $p_f$, which will decrease the normal stress $\sigma_n$ to the effective $\sigma_n^*$ [Hubbert and Rubey, 1959, Chapple, 1978]. The analysis of Sibson [1985],
which considers only the role of abnormally high $p_f$, without taking into account the role of the three other mechanisms, distinguishes three classes of fractures: (1) surfaces *favorably oriented* with respect to the stress axes, which can be reactivated with a hydrostatic $p_f$ for given differential stresses (see oblique dashed black lines and point A, Figure 3); (2) surfaces *moderately misoriented* with respect to the stress axes that will require large differential stresses or suprahydrostatic $p_f$ to be reactivated; (3) *severely misoriented* surfaces that will require, to be reactivated, $p_f$ values larger than in cases (1) or (2), with $\sigma_3^* < 0$ (see oblique black lines on Figure 3). The maximum differential stress allowing reactivation of severely misoriented surfaces is limited by the failure envelope. If the differential stress gets too high, a new and more favorably oriented shear surface or a vein will form instead of reactivating the severely misoriented surface (points B and C on Figure 3). In all cases, for a given value of differential stress, the $p_f$ value required to reactivate a severely misoriented surface is higher than that necessary for reactivation of a favorably oriented surface. The tectonic regime also has an impact on the $\lambda$ values for reactivation and vein formation. Indeed, reverse tectonic regimes generally require higher $p_f$ than normal stress regimes do (Figure 3).

## 2.5. The fault valve model

The *fault-valve* model, which is a consequence of the stress analysis of Sibson [1985], is based on the existence, in faulted regions, of fault planes moderately or severely misoriented with respect to the stress axes at the time of the fault activity [Sibson et al., 1988, Sibson, 1989, 1990]. The model is based on the following assumptions: (1) the coefficient of friction on the pre-existing (misoriented) fault is of “Byerlee” type, that is, its value is between 0.6 and 0.85; (2) the cohesion C of the surface is neglected; (3) the fluid does not chemically react with the rock, that is, only the fluid pressure $p_f$ plays a role.

The typical fault-valve behavior of a fault or fault system involves cyclical variations of $p_f$. A cycle consists of four stages: (1) $p_f$ increases up to suprahydrostatic values below a hydraulic barrier, (2) rupture of the fault plane following a Mohr-Coulomb type criterion [Sibson, 1985], (3) fluid escape (upward or sideward) following increase of the rock

permeability by fracturing and cataclasis associated with fault and hydraulic barrier ruptures, (4) fracture or pore sealing by mineral precipitation due to the sudden decrease of $p_f$ down to hydrostatic values. This last stage eventually results in a restauration of the former seal, thus bringing back the faulted rock mass to a state before stage (1), allowing a repetition of the cycle. Following the initial discovery of this mechanism in mesothermal gold-quartz deposits by Sibson et al. [1988], this behavior was recognized in ancient fault systems [Cox, 1995, Robert et al., 1995, Hacker, 1997, Nguyen et al., 1998, Sibson and Scott, 1998, Faleiros et al., 2007]. The extensive damage to the rocks adjacent to the faults suggests that reactivation might have been seismogenic. The fault-valve model is proposed to account for present-day seismic reactivation of unfavourably oriented faults (see Section 4).

## 2.6. The seismic pumping mechanism

Seismic pumping provides a mechanism of formation of zoned mineralized veins along faults in the brittle crust by calling for fluid flow during earthquakes [Sibson et al., 1975, Kerrich et al., 1987]. The mineral zonation is interpreted as the result of repeated arrivals of fluids, each arrival leaving an imprint in the vein. In the pre-seismic stage, due to tectonic loading along the fault, dilatant structures such as tension cracks will create voids along and near the fault, aspirating nearby fluids. In the post-seismic stage, fluids will be expelled upwards or sidwards, before the cycle starts again. In itself, the seismic pumping mechanism does not increase directly the $p_f$ value, but contributes to it by injecting volumes of fluids from remote areas towards active fault zones. It can also be active during diagenesis in faulted sedimentary basins [Wood and Boles, 1991].

## 2.7. Shear stress-enhanced compaction

In most cases, the central part, or core, of a fault zone is constituted by a weak material (clayey gouge, poorly consolidated breccia and so on). When tectonically loaded, a core material may undergo more compaction than the stronger surrounding rocks (damage zone, country rocks). If the core zone is (i) saturated by fluids and (ii) sealed (i.e., the fluid
cannot escape), then tectonic load-induced compaction will increase $p_f$ [Byerlee, 1990, Blanpied et al., 1992, Sleep and Blanpied, 1992]. The $p_f$ increase will then decrease the effective normal stress $\sigma_n^*$, therefore allowing or favoring fault slip. Note that this mechanism does not require any temperature increase.

## 2.8. Thermal pressurization

Thermal pressurization, also called “dynamic” pressurization, concerns thin (a few millimeters to a few centimeters) fault core zones which are the site of coseismic slip along faults in the upper crust [Sibson, 1973, Lachenbruch, 1980, Mase and Smith, 1987, Wibberley and Shimamoto, 2005]. The phenomenon is based on a contrast between the thermal expansion of aqueous fluids trapped in the very-low-permeability fault core rock (clayey gouge in most cases) and the thermal expansion of the fault core rock matrix itself. During co-seismic, slip, the fluid-saturated rock will thermally expand due to shear heating, but fluids will expand in larger proportions than the pores. As long as the rock permeability remains low (e.g., in the absence of fracturing), the fluid cannot escape from the core zone and, following expansion, its pressure will increase drastically, resulting in a decrease of the normal stress acting on the surface.

The recognition of thermal pressurization processes during earthquakes or in the geological record is a challenge and, in addition to observations of natural microstructures [Ujiie et al., 2007, 2010, Boullier et al., 2009, Boullier, 2011], must rely on experimental reproduction of seismic slip and on numerical modeling [Noda and Shimamoto, 2003, Segall and Rice, 2006, Ujiie et al., 2011, Boutareaud et al., 2008, Ferri et al., 2010, Kitajima et al., 2011, Viesca and Garagash, 2015]. In addition to thermal pressurization strictly speaking, frictional heating can also lead to thermal decomposition of minerals such as phyllosilicates or carbonates (calcite, siderite), resulting in the release of fluids (H₂O, CO₂) that can then be pressurized [Brantut et al., 2008, 2010, Hirono et al., 2008, Ferri et al., 2010, Jeanne et al., 2014]. Frictional melting of the fault surface during co-seismic slip in carbonate-rich rocks can also result in a release of large amounts of CO₂ following volatile exsolution from the melt [Famin et al., 2008]. In the thermal

decomposition or exsolution processes, the released fluids are considered to be pressurized and can consequently decrease the normal stress acting on the faults. Some authors contend that so-called injected cataclasites or injected gouges may result from co-seismic thermal pressurization [Lin et al., 2013, Lin, 2019].

# 3. Role of abnormally high $p_f$ in crustal tectonics

## 3.1. Role of abnormally high $p_f$ in thrust tectonics

In areas dominated by thin-skinned tectonics, thrust faults can accommodate significant (>10 km) horizontal displacements. To overcome the mechanical resistance of the allochtonous units to displacement, two mechanisms among the ones described above are frequently called for. The first mechanism (the translation made possible by weak ductile décollement layers) allows to account for the formation of several foreland thrust-and-fold belts such as the French-Swiss Jura belt or the Iranian Zagros belt [Jordan, 1992, Sherkati et al., 2006, Sommaruga et al., 2017, Lacombe and Mouthereau, 2002].

The other mechanism, that is, the translation of thrust sheets triggered or favored by supra-hydrostatic $p_f$, is inferred or ascertained in various tectonic settings. Table 1 provides examples of studies that call for supra-hydrostatic $p_f$ to account for horizontal displacements along thrusts in collision belts or foreland fold-and-thrust belts. The supra-hydrostatic $p_f$ lowers $\sigma_n^*$, thereby reducing the frictional resistance to motion and allowing a smooth displacement of the allochton units over their relative autochthonous basement. Evidence for supra-hydrostatic $p_f$ either come from direct methods, namely $p_f$ measurements in exploration boreholes, or from indirect methods, such as analyses of geological structures (e.g., dilatant crack-seal veins or implosion breccias) or modeling (critical Coulomb wedge, Coulomb stress change). Fluid inclusions preserved in mineralized veins can provide estimates of $p_f$ at the time of vein formation. As such, they are a powerful and reliable method for $p_f$ estimates.

The role of supra-hydrostatic $p_f$ build-up appears critical or ubiquitous in accretionary prisms, whatever active or inactive [Moore and Vrolijk, 1992].
<table><caption>**Table 1.** Examples of geological or geophysical evidence for displacement along low-angle reverse faults triggered or favored by supra-hydrostatic pore fluid pressures *p*<sub>f</sub> inside or near fault zones in collision belts or foreland fold-and-thrust belts, with pore fluid ratio λ estimates where available</caption><thead><tr><th>Area</th><th>Name of structure</th><th>Evidence</th></tr></thead><tbody><tr><td>Central Alps</td><td>Glarus thrust</td><td>Fracturing, veining, brecciation alternating with ductile deformation (λ > 1)<sup>[1]–[4]</sup></td></tr><tr><td rowspan="2">Western Alps</td><td>Helvetic Diablerets nappe basal thrust</td><td>Large amounts of connate and metamorphic waters expelled upward, as indicated by δ<sup>18</sup>O, δ<sup>13</sup>C and δD in calcite or quartz veins, possibly overpressured by rapid tectonic burial<sup>[5]</sup></td></tr><tr><td>Sub-Alpine frontal thrust zone (Chartreuse and Vercors) and foreland (Valence basin)</td><td><i>p</i><sub>f</sub> measurements in exploration boreholes<sup>[6]</sup></td></tr><tr><td>Pyrenees</td><td>Gavarnie thrust and splay faults</td><td>Fluid inclusions in quartz veins indicate fluctuations of <i>p</i><sub>f</sub> between lithostatic (~500 MPa) and hydrostatic (~200 MPa)<sup>[7]</sup>; combination of shear reactivation and extension vein opening (λ = 0.77–0.89)<sup>[8]</sup></td></tr><tr><td>Himalaya</td><td>Himalayan Main Central Thrust</td><td>Critical Coulomb wedge modeling suggests λ = 0.8–0.9<sup>[9]</sup>, Coulomb stress change modeling suggests λ = 0.3–0.9<sup>[10]</sup></td></tr><tr><td>Apennines</td><td>Apenninic basal thrust zone, Apenninic foreland</td><td>Fluid inclusions in quartz veins indicate near-lithostatic <i>p</i><sub>f</sub> at the time of the emplacement of the Liguride nappes<sup>[11]</sup>; calcite-cemented implosion breccias filling dilational jogs (<i>p</i><sub>f</sub> > σ<sub>3</sub>)<sup>[12]</sup></td></tr><tr><td>Taiwan</td><td>Taiwan fold-and-thrust belt basal decollement fault, Chelungpu fault</td><td><i>p</i><sub>f</sub> measurements in exploration boreholes (λ ~ 0.7)<sup>[13]</sup>, critical Coulomb wedge modeling<sup>[14],[15]</sup>; discrete hydraulic fractures<sup>[16]</sup></td></tr><tr><td>Western California</td><td>Coast Ranges and western Great Valley</td><td><i>p</i><sub>f</sub> measurements in exploration boreholes<sup>[17],[18]</sup></td></tr></tbody></table>

References: [1]Burkhard et al. [1992], [2]Badertscher and Burkhard [2000], [3]Badertscher et al. [2002], [4]Hürzeler and Abart [2008], [5]Crespo-Blanc et al. [1995], [6]Deville [2021], [7]Henderson and McCaig [1996], [8]Lacroix et al. [2013], [9]Mugnier et al. [1994], [10]Cattin and Avouac [2000], [11]Mullis [1988], [12]Vannucchi et al. [2010], [13]Suppe and Wittke [1977], [14]Davis et al. [1983], [15]Barr and Dahlen [1990], [16]Boullier [2011], [17]Melchiorre et al. [1999], [18]Unruh et al. [1992].

Table 2 provides examples of various continental margins or paleo-margins in which displacement along past or present (first-order) décollement faults or second-order faults (also referred to as splay faults) branching on décollements was or is possible due to supra-hydrostatic *p*<sub>f</sub>. Similarly to what is inferred along large-displacement thrusts, the supra-hydrostatic *p*<sub>f</sub> lowers *σ*<sub>n</sub><sup>*</sup>, thereby reducing the frictional resistance to motion and allowing a smooth

displacement of the overriding plate oceanwards, at least in the shallowest parts of the plate interface, before pressure and temperature increases lead to dehydration and expulsion of fluids. Direct evidence for supra-hydrostatic *p*<sub>f</sub> along the basal décollements of accretionary prisms come from *p*<sub>f</sub> measurements in Deep-Sea Drilling Project (DSDP), Ocean Drilling Program (ODP) or Integrated Ocean Drilling Program (IODP) boreholes. Porosity measurements
<table><caption>**Table 2.** Examples of geological or geophysical evidence for supra-hydrostatic pore fluid pressures *p*<sub>f</sub> measured or inferred along low-angle reverse faults in accretionary prisms</caption><thead><tr><th>Type of prism</th><th>Area</th><th>Type of structure</th><th>Evidence</th></tr></thead><tbody><tr><td rowspan="4">Active accretionary prisms</td><td>Cascadia margin</td><td>Frontal</td><td rowspan="2">Seismic velocity variations and negative polarity reflections in seismic profiles<sup>[11–13]</sup>; ODP core consolidation tests and logging data (λ = 0.88–0.99)<sup>[4]</sup>; supra-hydrostatic p<sub>f</sub> (λ = 0.8) computed from velocity-depth data<sup>[5]</sup>; reverse polarity reflections from the décollement<sup>[6]</sup></td></tr><tr><td>Oregon accretionary prism</td><td>décollement</td></tr><tr><td>Central America Costa Rica and Guatemala margins</td><td>Frontal</td><td rowspan="2">Direct measurement of p<sub>f</sub> in DSDP boreholes (λ = 0.25–0.48)<sup>[7]</sup>; consolidation tests on ODP cores suggest excess p<sub>f</sub> (1.3–3.1 MPa)<sup>[8]</sup>; numerical modeling of p<sub>f</sub> build-up following sediment dewatering in undrained conditions<sup>[9]</sup>; p<sub>f</sub> measurements in ODP borehole hydrologic observatories (λ = 0.25–1)<sup>[10]</sup>; pervasive veining suggests p<sub>f</sub> fluctuations in accordance with the seismic cycle<sup>[11]</sup></td></tr><tr><td>SW Japan Nankai</td><td>Frontal</td></tr><tr><td>Trough accretionary prism</td><td>décollement</td><td>Shear tests on saturated clay samples from the décollement indicate that friction drops dramatically with increasing p<sub>f</sub><sup>[12]</sup>; overconsolidation and dilatant fractures along the décollement<sup>[13]</sup>; seismic reflection amplitude variations along the décollement<sup>[14]</sup>; modeling of p<sub>f</sub> based on acoustic velocity logs<sup>[15]</sup>; modeling of p<sub>f</sub> based on laboratory permeability measurements from core samples suggest λ* = 0.68–0.77<sup>[16]</sup>; supra-hydrostatic p<sub>f</sub> deduced from seismic interval velocities<sup>[17]</sup>; empirical relationships between V<sub>P</sub>, porosity, and effective σ<sub>m</sub> obtained from laboratory deformation tests on drill core samples indicate excess p<sub>f</sub> of 17–87 MPa<sup>[18]</sup>; upward flow of drilling mud from IODP borehole indicates an initial excess p<sub>f</sub> of ~5–10 MPa above p<sub>H</sub><sup>[19]</sup>; estimation of p<sub>f</sub> from empirical relationships between p<sub>f</sub>, porosity and V<sub>P</sub>, obtained from consolidation experiments and from drilling or sonic velocity data<sup>[20]</sup></td></tr><tr><td>Lesser Antilles</td><td>Frontal</td><td rowspan="2">High-amplitude negative polarity reflections in seismic profiles<sup>[21], [22]</sup>; low impedance layer<sup>[23]</sup></td></tr><tr><td>Barbados Ridge accretionary prism</td><td>décollement</td></tr><tr><td>New Zealand</td><td>Papaku splay fault</td><td>p<sub>f</sub> estimated from seismic velocity or from negative polarity reflections<sup>[24], [25]</sup>; direct measurements in ODP boreholes and sediment consolidation tests (λ up to >0.9)<sup>[26], [27]</sup>; high p<sub>f</sub> suggested by anisotropy of magnetic susceptibility<sup>[28]</sup></td></tr><tr><td>Hikurangi accretionary prism</td><td></td><td>Rock deformation experiments on cores from the fault zone (λ* = 0.3–0.6)<sup>[29]</sup></td></tr></tbody></table>

(continued on next page)
<table><caption>**Table 2.** (continued)</caption><thead><tr><th>Type of prism</th><th>Area</th><th>Type of structure</th><th>Evidence</th></tr></thead><tbody><tr><td rowspan="2">Exhumed inactive accretionary prisms</td><td rowspan="2">Shimanto accretionary prism (Japan)</td><td>Nobeoka paleo-décollement</td><td rowspan="2">Fluid inclusions and modeling suggest p<sub>f</sub> excess values between 5 and 20 MPa<sup>[30]</sup></td></tr><tr><td>Nakanomata splay fault</td></tr><tr><td rowspan="2">Kodiak accretionary prism (Alaska)</td><td rowspan="2">Paleo-décollements and mélange zones</td><td>Crack-seal quartz veins parallel to ancient décollements interpreted as resulting from repeated hydrofracturing (p<sub>f</sub> > σ<sub>3</sub>)<sup>[32]</sup>; fluid inclusions indicate near-lithostatic p<sub>f</sub> (p<sub>f</sub> ~ σ<sub>1</sub>) at the time of quartz vein formation<sup>[33],[34]</sup>; lack of low-friction material and mechanical analysis of displacement along the décollement plead for supra-hydrostatic p<sub>f</sub><sup>[35]</sup>; hydrofractures suggest supra-hydrostatic p<sub>f</sub><sup>[36]</sup>; fluid inclusions in hydrofractures suggest near-lithostatic p<sub>f</sub><sup>[37]</sup></td></tr><tr><td>Mudstone matrix of mélange</td><td>Shear veins severely misoriented with respect to stress field<sup>[38],[39]</sup></td></tr><tr><td>Chrystalls Beach accretionary prism (New Zealand)</td><td></td><td></td><td></td></tr></tbody></table>

References: [1] Cochrane et al. [1994], [2] Cochrane et al. [1996], [3] Moore et al. [1995a], [4] Moore et al. [1995b], [5] Hyndman et al. [1993], [6] Tobin et al. [1994], [7] Von Huene [1985], [8] Saffer et al. [2000], [9] Screaton and Saffer [2005], [10] Davis and Villinger [2006], [11] Vannucchi and Leoni [2007], [12] Brown et al. [2003], [13] Ujijie et al. [2003], [14] Bangs et al. [2004], [15] Tsuji et al. [2008], [16] Skarbek and Saffer [2009], [17] Tobin and Saffer [2009], [18] Kitajima and Saffer [2012], [19] Hirose et al. [2021], [20] Pwavodi and Doan [2024], [21] Park et al. [2000], [22] Park et al. [2002], [23] Bangs et al. [1990], [24] Shipley et al. [1994], [25] Moore et al. [1995c], [26] Moore and Tobin [1997], [27] Moore et al. [1998], [28] Housen et al. [1996], [29] French and Morgan [2020], [30] Raimbourg et al. [2015], [31] Passelegue et al. [2014], [32] Fisher and Byrne [1987], [33] Vrolijk [1987], [34] Vrolijk et al. [1988], [35] Byrne and Fisher [1990], [36] Fisher et al. [1995], [37] Rowe et al. [2009], [38] Fagereng et al. [2010], [39] Fagereng et al. [2011].

ODP: Ocean Drilling Program. IODP: International Ocean Drilling Program. 3D: three-dimensional. p<sub>H</sub>: hydrostatic pressure. σ<sub>m</sub>: mean stress (σ<sub>m</sub> = (σ<sub>1</sub> + σ<sub>2</sub> + σ<sub>3</sub>)/3). λ*: modified pore fluid factor, λ* = (p<sub>F</sub> - p<sub>H</sub>)/(p<sub>L</sub> - p<sub>H</sub>), where p<sub>H</sub> is the hydrostatic pressure and p<sub>L</sub> the lithostatic pressure caused by the weight of the overburden [Saffer and Tobin, 2011].
and various mechanical tests (e.g., consolidation tests) on DSDP/ODP/IODP cores further testify for under-compaction of sediments or rocks along the décollements, accounted for by the assumption of supra-hydrostatic to nearly lithostatic $p_f$. The presence of hydrofractures along as well as above or below décollements is also direct evidence [Table 2; see also Behrmann, 1991]. Indirect evidence come from seismic reflection imaging, with characteristic negative polarity reflections along décollement surfaces. Lastly, the weakening role of supra-hydrostatically pressurized fluids along the basal décollement of accretionary prisms can be reproduced by analog modelling [Cobbold et al., 2001, Mourgues and Cobbold, 2003].

## 3.2. Role of abnormally high $p_f$ in extensional tectonics

Large displacement along low-angle normal faults, also referred to as detachments, is mechanically paradoxical. The different mechanisms mentioned above to account for the displacement along misoriented fault surfaces apply to low-angle normal faults [Reynolds and Lister, 1987, Yin, 1989, Axen, 1992, Westaway, 1999, Abers, 2009, Collettini, 2011]. Table 3 provides examples of studies which suggest or demonstrate that reactivation of low-angle normal faults was possible because of supra-hydrostatic to lithostatic $p_f$ values. Regarding inactive detachments, structural evidence for high $p_f$ come from textures of precipitated minerals in breccia cements (so-called cockade structures) or in flat-lying crack-seal dilatant veins. Fluid inclusions in the precipitated minerals also provide indications. Lastly, mechanical modeling also accounts for overpressured-fluid-aided displacements along low-angle faults. Regarding active detachments, evidence consist of vein or breccia cement texture analyses (including fluid inclusions), mechanical modeling, $p_f$ measurements in boreholes, geophysical investigations (body wave velocity anomalies).

The other mechanisms listed above may also facilitate large displacement along low-angle normal faults. In particular, analog modeling suggests that extension can take place above continuous low-strength layers such as evaporites [Vendeville and Jackson, 1992a,b]. Regarding the reorientation of the stress axes in the vicinity of the fault, Lecomte

et al. [2011] propose to explain reactivation of misoriented low-angle normal faults by calling for an elasto-plastic frictional gouge instead of the classical frictional fault gouge. Plastic strain can initiate the displacement despite the misorientation. Following the initiation, stress axes near the fault zone rotate to a more favorable orientation with respect to the fault surface, allowing further displacement.

## 3.3. Role of abnormally high $p_f$ in strike-slip tectonics

The activity of steeply dipping to vertical plate-boundary transform or intra-plate strike-slip faults moderately to severely misoriented with respect to the principal stress axes constitutes a challenge since the discovery of nearly perpendicular normal stresses acting along the San Andreas fault [Mount and Suppe, 1987, Zoback et al., 1987]. As such, the San Andreas faults appears as a weak fault, and explanations for it have been extensively looked for. The models and mechanical analyses of Byerlee [1990, 1993] and Rice [1992] attempt to show that supra-hydrostatic $p_f$ could account for the weakness of the fault. Experimental measurement of clay permeabilities [e.g., Faulkner and Rutter, 2001] provide a mechanism for the maintenance of excess $p_f$ in strike-slip fault zones. However, drilling across the fault zone during the SAFOD (San Andreas Fault Observatory at Depth) program at ~3 km depth did not reveal any supra-hydrostatic $p_f$, prompting researchers to propose other mechanisms that could explain the weakness: fault rocks with extremely low friction coefficients such as talc [Moore and Rymer, 2007] or clays [Schleicher et al., 2009a,b, Carpenter et al., 2011, Holdsworth et al., 2011], stress axes reorientations to more favorable attitudes at short distances from the fault [Provost and Houston, 2001, Hardebeck and Michael, 2004, Healy, 2008]. It seems clear that pressurized fluids do not weaken the San Andreas fault in its upper part (between the surface and a ~3 km depth), except perhaps for micro-earthquake activity [Mittempergher et al., 2011]. Hydro-mechanical modeling [e.g., Fulton and Saffer, 2009, Beeler et al., 2013] suggests that supra-hydrostatic $p_f$ could play a role in the weakening at depths larger than the SAFOD drilling. Conversely, hydro-mechanical modeling by Fulton et al. [2009] suggest that supra-hydrostatic $p_f$ cannot
<table><caption>**Table 3.** Examples of geological or geophysical evidence for displacement along low-angle normal faults (detachment faults) triggered or favored by supra-hydrostatic pore fluid pressures $p_f$ inside or near the fault zone</caption><thead><tr><th>Type of low-angle normal fault</th><th>Name of structure, area</th><th>Evidence</th></tr></thead><tbody><tr><td rowspan="4">Inactive intra-continental</td><td>Western US Basin and Range (Southern Mountains MCC deformed zones, Arizona; Whipple detachment, California; Siever Desert detachment, Utah)</td><td>Flat-lying dilatant veins suggest $p_f > \sigma_3^{[1]}$; fluid inclusions in quartz veins suggest $p_f \ge 120$ MPa<sup>[2],[3]</sup>; critical Coulomb wedge analysis suggests $p_f \sim 0.6 \sigma_v^{[4]}$</td></tr><tr><td>Central Alps (Brenner, Simplon and Grimsel faults)</td><td>Fluid inclusions in veins suggest supra-hydrostatic $p_f^{[5]}$; mylonitic zone embrittlement suggests episodic supra-hydrostatic $p_f^{[6],[7]}$; textural characteristics of cockade structures in a fault breccia and mechano-chemical calculations allow to reconstruct $p_f$ fluctuations related to different stages of the seismic cycle<sup>[8]</sup></td></tr><tr><td>Gulf of Corinth LANF, Greece</td><td>Limit Analysis (maximum strength theorem analysis) suggests $p_f = 0.57-0.77 \sigma_v^{[9]}$</td></tr><tr><td>North Apennines LANFs, including Alto Tiberina and Tellaro detachments</td><td>Limit Analysis (maximum strength theorem analysis) suggests that $p_f = 0.57-0.77 \sigma_v^{[9]}$; $p_f \sim 0.85 \sigma_v$ in exploration boreholes<sup>[10]</sup>; brittle tensile structures and fluid inclusions in calcite veins suggest episodic supra-hydrostatic $p_f^{[11]}$; $V_P$ positive anomaly, $V_P/V_S$ negative anomaly<sup>[12]</sup></td></tr><tr><td>Inactive intra-oceanic or intra-OCT</td><td>Detachment surfaces exposed in Alpine ophiolites</td><td>Calcite-cemented breccias (ophicalcites)<sup>[13],[14]</sup>; gouge veins injected in cataclasite suggest transient supra-hydrostatic $p_f^{[15]}$</td></tr><tr><td>Active intra-oceanic</td><td>Mid-Atlantic ridge Atlantis massif, mid-Atlantic ridge 13°–15° N<br>South Pacific Woodlark-D'Entrecasteaux ridge Moresby detachment</td><td>Fluid inclusions<sup>[16],[17]</sup>; hydraulic breccias<sup>[18]</sup><br>Low $V_P$ and high porosity values in the fault zone suggest supra-hydrostatic $p_f^{[19]}$; dilatant crak-seal veins suggest episodic near-lithostatic $p_f$ and hydrofracturing<sup>[20],[21]</sup>; calcite veins parallel to foliation suggest supra-lithostatic $p_f^{[22]}$</td></tr></tbody></table>

References: [1] Reynolds and Lister [1987], [2] Smith et al. [1991], [3] Selverstone et al. [2012], [4] Yuan et al. [2018], [5] Selverstone et al. [1995], [6] Axen et al. [1995], [7] Axen et al. [2001], [8] Berger and Herwegh [2019], [9] Yuan et al. [2020], [10] Collettini et al. [2008], [11] Clemenzi et al. [2015], [12] Moretti et al. [2009], [13] Früh-Green et al. [1990], [14] Picazo et al. [2013], [15] Manatschal [1999], [16] Escartin et al. [2003], [17] Castelain et al. [2014], [18] Picazo et al. [2012], [19] Floyd et al. [2001], [20] Roller et al. [2001], [21] Kopf et al. [2003], [22] Famin and Nakashima [2005].

MCC: metamorphic core complex. LANF: low-angle normal fault. OCT: ocean-continent transition. $\sigma_v$: principal vertical stress axis.
be maintained along the San Andreas fault at shallow (<3 km) depth, and that other mechanisms such as shear-enhanced compaction or thermal pressurization inside the fault zone have to be invoked.

Table 4 provides examples of transform or strike-slip faults along which supra-hydrostatic $p_f$ could play or could have played a role in the past or present fault activity. Direct evidence for abnormally high $p_f$ can be found in shallow boreholes drilled across the New Zealand Alpine fault, but no data from depths larger than a few hundred meters are available. Indirect evidence consist either in rare hydraulic fractures or in more frequently observed geophysical (electrical conductivity or seismic velocity) anomalies. In this last case, even if the anomalies can be interpreted as caused by the presence of fluids at depth, they do not bring indisputable evidence for pressurized fluids. The apparent scarcity of abnormally high $p_f$ ascertained along transform or strike-slip fault zones may be due to the lack or scarcity of efficient seals. Indeed, fault zones, especially their damage zones if any, crossing the entire crust may appear as conduits rather than barriers or seals.

### 3.4. Conclusions

A review of the literature dealing with the role of abnormally high $p_f$ along faults in the brittle crust allows to draw the following key ideas.

(1) Undebatable evidence for supra-hydrostatic $p_f$ come from pressure measurements in oil industry or scientific boreholes (DSDP, OPD, IODP). Near-horizontal dilatant veins found along low-angle faults also constitute a strong evidence, especially when fluid inclusions preserved in these veins indicate supra-hydrostatic $p_f$ at the time of fluid entrapment. Geophysical anomalies, either electrical (conductivities) or seismological (seismic velocities), are interpreted as caused by abnormally high $p_f$ at depth along fault zones.

(2) Evidence, whatever direct or indirect, for supra-hydrostatic $p_f$ appear common along low-angle normal faults (detachments) or reverse faults (including décollements) (Tables 1–3). Fluids appear to be efficiently trapped beneath flat-lying or gently dipping allochtons that can act as impervious or low-permeability lids. Conversely, conclusive evidence for supra-hydrostatic $p_f$ along transform or strike-slip faults are scarcer than along low-angle

normal or reverse faults (Table 4), possibly because such vertical or steeplydipping fault zones lack tight and laterally continuous seals that could prevent fluids from leaking upward to the surface.

(3) Along with fracturing and comminution linked with displacement along faults, fluid circulation and associated fluid-rock interactions quite often result in alteration of minerals, resulting in weaker minerals such as clays, talc and other minerals whose frictional properties are much weaker than primary minerals. This chemical effect of fluids is clearly a serious competitor with the pressurization of fluids to account for fault weakening.

(4) Mechanisms such as shear stress-enhanced compaction, thermal pressurization, or stress reorientation near fault zones certainly play roles in the reactivation of misoriented faults, but they are difficult to be recognized in the geological record.

## 4. Fluids and natural earthquakes

### 4.1. Intra-plate earthquakes, earthquake sequences, background seismicity, subduction earthquakes

Table 5 presents examples of studies that attempt to relate major intra-continental earthquakes, earthquake sequences, background seismicity, subduction earthquakes and aftershock sequences to abnormally high $p_f$. Some of these examples were already introduced in the previous section, but without any specific mention to the seismogenic activity of the relevant faults.

Except rare examples of direct measurements in boreholes, evidence for supra-hydrostatic $p_f$ triggering or favoring intra-plate earthquakes are mostly indirect. Body-wave velocity anomalies detected by two-dimensional or three-dimensional tomography suggest that pressurized fluids are present in hypocentral regions of many earthquakes. Most authors agree that fluids are trapped at depth and that their accumulation results in abnormally high $p_f$ values which in turn weaken the faults.

Evidence for abnormally high $p_f$ that could either trigger or favor the propagation of subduction zone earthquakes mainly comes from body wave velocity anomalies revealed by tomography (Table 5). Other indirect evidence come from mechanical analyses based on the reactivation of misoriented fault
<table><caption>**Table 4.** Examples of geological or geophysical evidence for displacement along transform faults or strike-slip faults triggered or favored by supra-hydrostatic pore fluid pressures *p*<sub>f</sub> inside or near fault zones</caption><thead><tr><th>Area, name of structure</th><th>Evidence</th></tr></thead><tbody><tr><td>California, San Andreas fault</td><td>Very low resistivity, low V<sub>P</sub>, and high a V<sub>P</sub>/V<sub>S</sub> ratio suggest supra-hydrostatic p<sub>f</sub><sup>[1]</sup>; calcite and anhydrite-filled mode I veins in SAFOD cores suggest that transient increases of supra-hydrostatic (possibly supra-lithostatic) p<sub>f</sub> occurred during fault activity and may constitute a triggering mechanism for some micro-earthquakes recorded at depth near SAFOD drilling site<sup>[2]</sup>; frictional sliding and stress-driven DMT microstructures suggest that deformation in the active shear zones is displacement-weakening, possibly due to local and transient high p<sub>f</sub> build-ups along creeping segments<sup>[3]</sup>; microstructural and geochemical data from SAFOD samples indicate that transient co-seismic fluid overpressure events overprint aseismic creep along the fault<sup>[4]</sup></td></tr><tr><td>Turkey, North Anatolian Fault</td><td>Fluid inclusions in subhorizontal extension veins suggest supra-hydrostatic p<sub>f</sub><sup>[5]</sup>; continuous ascent of deep crustal or mantle fluid (CO<sub>2</sub>, CH<sub>4</sub> or He) could result in near-lithostatic or supra-lithostatic p<sub>f</sub><sup>[6]</sup>; low electrical resistivity (measured by magnetotelluric methods) and low V<sub>P</sub> (measured by 2-D seismic velocity tomography) suggest high p<sub>f</sub> at 5–15 km depth<sup>[7],[8],[9]</sup></td></tr><tr><td>New Zealand, Alpine Fault</td><td>Slug tests in shallow (150 m deep) boreholes suggest that the near-surface Alpine fault zone may be site of high p<sub>f</sub> gradient<sup>[10]</sup>; pressure measurement during drilling reveals a slight (9%) supra-hydrostatic p<sub>f</sub> in shallow (~900 m deep) borehole<sup>[11]</sup>; tremors appear to be located at 25–45 km depth, in a region of high P-wave attenuation that could be the site of supra-hydrostatic p<sub>f</sub><sup>[12]</sup></td></tr><tr><td>Japan, Atotsugawa fault</td><td>High p<sub>f</sub> could account for creep along some parts of the fault<sup>[13]</sup></td></tr><tr><td>Japan, Nojima fault</td><td>Discrete hydraulic fractures<sup>[14]</sup></td></tr><tr><td>Central Alps, Adamello pluton</td><td>Reactivation of the severely misoriented inactive Gole Larghe fault<sup>[15]</sup></td></tr><tr><td>Southeastern Brazil, Ribeira Shear Zone</td><td>Fluid inclusions in foliation-parallel or extensional quartz veins along the shear zone indicate p<sub>f</sub> fluctuations that are interpreted as the result of a fault-valve behavior<sup>[16]</sup></td></tr></tbody></table>

References: [1]Eberhart-Phillips and Michael [1993], [2]Mittempergher et al. [2011], [3]Hadizadeh et al. [2012], [4]Hadizadeh et al. [2024], [5]Janssen et al. [1997], [6]Pfister et al. [2000], [7]Karabulut et al. [2003], [8]Tank et al. [2005], [9]Karaş et al. [2020], [10]Sutherland et al. [2012], [11]Sutherland et al. [2017], [12]Wallace [2020], [13]Kato et al. [2007], [14]Boullier [2011], [15]Mittempergher et al. [2009], [16]Faleiros et al. [2007].

SAFOD: San Andreas Fault Observatory at Depth. DMT: diffusive mass transfert.
<table><caption>**Table 5.** Examples of geophysical evidence for major earthquakes and aftershocks triggered or favored by abnormally high pore fluid pressures $p_f$ in or around fault zones</caption><thead><tr><th>Type of event</th><th>Area and/or events</th><th>Evidence</th></tr></thead><tbody><tr><td>Intraplate earthquakes, earthquake sequences and background seismicity</td><td>Central Alps<br>Himalaya MCT<br>Himalaya background sesimicity<br>Northern Apennines (1997–1998 Umbria-Marche-Colfiorito, 2000 Faenza and 2012 Emilia seismic sequences)<br>Central and southern Apennines (2009 Abruzzo-L'Aquila sequence, 2010–2014 Mt Pollino sequence)</td><td>Unexpected background seismic activity in the lower crust accounted for by supra-hydrostatic $p_f^{[1]}$<br>Drop in $V_p$ and $V_S$ values, increase of $V_p/V_S^{[2]}$<br>Supra-hydrostatic $p_f$ at depth deduced from correlated seasonal variation of stress and seismicity<sup>[3]</sup><br>Supra-hydrostatic $p_f$ ($\lambda \ge 0.9$) inferred from the presence at depth of Triassic evaporite impervious seals<sup>[4]</sup>; supra-hydrostatic $p_f$ ($\lambda \ge 0.7$) suggested by stress inversion of focal mechanisms<sup>[5]</sup>; space-time evolution of seismic activity fits with diffusion laws, $p_f \sim \sigma^{[6]}$<sup>[7]</sup>; $V_p/V_S$ changes along a fault system suggest that $p_f$ fluctuations controlled the space-time evolution of the Emilia sequence and the activation of the second main shock<sup>[8]</sup><br>Supra-hydrostatic CO<sub>2</sub> $p_f$ values in exploratory boreholes<sup>[9], [10]</sup>; supra-hydrostatic $p_f$ propagation along a fault system tracked by space-time variations of $V_p/V_S$ anomalies during earthquake migration<sup>[11]</sup>; high $V_p/V_S$ and low $Q_p/Q_S$ indicative of supra-hydrostatic $p_f^{[12]}$; temporal evolution of $V_p/V_S$ suggests an increase through time of supra-hydrostatic $p_f$ (up to ~200 MPa at 10 km depth), eventually triggering the 6 April 2009 M<sub>w</sub> 6.3 mainshock<sup>[13]</sup>; increase of the $V_p/V_S$ ratio and evolution of shear wave splitting before the earthquake are interpreted as reflecting the progressive influx of pressurized fluids near the hypocentral of the M<sub>w</sub> 6.3 L'Aquila main shock<sup>[14]</sup>; mantle-sourced CO<sub>2</sub> accumulates under near-lithostatic or supra-lithostatic $p_f$ in deep traps which may control the Apennines seismic activity<sup>[15], [16]</sup></td></tr><tr><td>East Iceland rift zone</td><td></td><td>Stress inversion of focal mechanisms suggests near-lithostatic $p_f$ values ($\lambda \sim 1$)<sup>[17]</sup></td></tr><tr><td>Back-arc side of NE Japan ($M > 6$ thrust-fault type events)</td><td></td><td>Mohr-Coulomb analysis suggests that severely misoriented faults are reactivated by supra-hydrostatic $p_f$ through a fault-valve mechanism<sup>[18]</sup></td></tr></tbody></table>

(continued on next page)
<table><caption>**Table 5. (continued)**</caption><thead><tr><th>Type of event</th><th>Area and/or events</th><th>Evidence</th></tr></thead><tbody><tr><td></td><td>Kaapvaal craton<br>(2017 M<sub>w</sub> 6.5<br>Bostwana<br>earthquake)</td><td>Episodic seismic activity in an otherwise stable craton explained by sporadic near-lithostatic p<sub>f</sub><br>(λ ~ 0.85)<sup>[19]</sup></td></tr><tr><td>New Madrid seismic<br>zone (1811–1812<br>sequence)</td><td></td><td>A test of the role of stress tensor shape Φ, fault friction coefficient μ, p<sub>f</sub>, and gradient of σ<sub>D</sub> with depth<br>by a Mohr–Coulomb-based parametric analysis shows that reactivation of the faults requires<br>supra-hydrostatic p<sub>f</sub> (λ = 0.68–0.81), unless μ is less than 0.4<sup>[20]</sup></td></tr><tr><td>Japan, Nojima fault,<br>1995 M<sub>w</sub> 7.2 Kobe<br>earthquake</td><td></td><td>Seismic velocity anomalies near the hypocenter could reflect supra-hydrostatic p<sub>f</sub> that may have<br>contributed to the initiation of the Kobe earthquake<sup>[21]</sup></td></tr><tr><td>Cascadia</td><td></td><td>Seismic tomography indicates abnormal V<sub>p</sub>/V<sub>s</sub> and Poisson ratio values interpreted as caused by<br>near-lithostatic p<sub>f</sub><sup>[22], [23]</sup></td></tr><tr><td>NE Japan</td><td></td><td>Aftershock focal mechanisms of the 1968 M 8.2 Tokachi earthquake suggest a weak plate interface<br>accounted for by the post-seismic release of overpressured fluids following the main rupture<sup>[24]</sup>; seismic<br>reflection imaging shows that the megathrust corresponds to a reflector with negative polarity, which<br>could be explained by high p<sub>f</sub> having possibly triggered the the 2011 M<sub>w</sub> 9 Tohoku event<sup>[25]</sup>;<br>mechanical analysis suggests stress permutations in the forearc region caused by fluctuations in p<sub>f</sub><br>during the seismic cycle (fault-valve action)<sup>[26]</sup></td></tr><tr><td>SW Japan</td><td></td><td>3-D seismic reflection data show the presence of a LVZ (~2 km thick, ~15 km wide, and ~120 km long)<br>beneath the Nankai accretionary prism, inside which inferred abnormally high p<sub>f</sub> may control<br>interplate coupling<sup>[27]</sup></td></tr><tr><td>Central America</td><td></td><td>Extremely high V<sub>p</sub>/V<sub>s</sub> ratios in the entire subducting oceanic crust is interpreted as due to high p<sub>f</sub>.<br>Varying velocity ratios in the overriding continental crust further allow to map varying coupling<br>domains<sup>[28]</sup></td></tr><tr><td>Chile</td><td></td><td>The importance of the post-seismic fluid discharge suggests overpressured p<sub>f</sub> beneath the subduction<br>interface before large interplate events<sup>[29], [30]</sup>; 3-D seismic tomography allows to map the distribution,<br>along the plate interface, of domains with abnormally high p<sub>f</sub> and domains with lower p<sub>f</sub> and to<br>propose a correlation between these domains and the patches which slip aseismically (high p<sub>f</sub>) and<br>those which are locked (low p<sub>f</sub>)<sup>[31]</sup></td></tr></tbody><tfoot><tr><td colspan="3" style="text-align: right;">(continued on next page)</td></tr></tfoot></table>
<table><caption>**Table 5.** (continued)</caption><thead><tr><th>Type of event</th><th>Area and/or events</th><th>Evidence</th></tr></thead><tbody><tr><td rowspan="5">Aftershocks</td><td>New Zealand, Hikurangi subduction zone</td><td>A Mohr–Coulomb analysis suggests that the recurrence of megathrust earthquakes may be triggered by fluctuation of p<sub>f</sub> (from hydrostatic to lithostatic) in addition to stress accumulation<sup>[32]</sup></td></tr><tr><td>1992 Landers earthquake</td><td>Increase of p<sub>f</sub> because of post-seismic fluid flow and Coulomb stress changes account for the chronology and spatial migration of the Landers event aftershocks<sup>[33]</sup></td></tr><tr><td>Apennines (1997 Umbria-Marche, 2009 L'Aquila, 2012 Emilia earthquake sequence)</td><td>Space–time migration of aftershocks fits with diffusion laws of fluids in porous media<sup>[34], [35]</sup>; inversion of earthquake nodal planes allows to map the 3-D p<sub>f</sub> distribution at depth in the hypocentral region of the L'Aquila earthquake and shows that time-space migration of aftershocks fits with diffusion laws of p<sub>f</sub><sup>[36]</sup>; aftershocks induced by a reduction in the fault shear strength due to a pulse of p<sub>f</sub><sup>[37]</sup>; aftershock decay rate and V<sub>P</sub>/V<sub>S</sub> changes along the fault system suggest that a p<sub>f</sub> pulse controlled the space–time evolution of the 2012 Emilia sequence<sup>[38]</sup></td></tr><tr><td>2011 Tohoku earthquake, NE Japan</td><td>Reactivation of misoriented fault planes of the aftershock seismic faults is accounted for by varying supra-hydrostatic p<sub>f</sub><sup>[39]</sup></td></tr><tr><td>2004 Sumatra earthquake</td><td>Aftershock migration along a splay fault merging with the megathrust is explained by upward migration of fluids, although supra-hydrostatic p<sub>f</sub> is not ascertained<sup>[40]</sup></td></tr></tbody></table>

References: [1] Deichmann [1992], [2] Mandal et al. [2022], [3] Bettinelli et al. [2008], [4] Quattrochi [1999], [5] Boncio and Bracone [2009], [6] Lombardi et al. [2010], [7] Calderoni et al. [2009], [8] Pezzo et al. [2018], [9] Collettini et al. [2009a], [10] Chiodini et al. [2004], [11] Chiarabba et al. [2009a], [12] Chiarabba et al. [2009b]), [13] Di Luccio et al. [2010], [14] Lucente et al. [2010], [15] Chiodini et al. [2011], [16] Di Luccio et al. [2022], [17] Plateaux et al. [2012], [18] Sibson [2007, 2009], [19] Gardonio et al. [2018], [20] Leclère and Calais [2019], [21] Zhao et al. [1996], [22] Audet et al. [2009], [23] Peacock et al. [2011], [24] Magee and Zoback [1993], [25] Kimura et al. [2012], [26] Sibson [2013], [27] Park et al. [2010], [28] Audet and Schwartz [2013], [29] Husen and Kissling [2001], [30] Koerner et al. [2004], [31] Moreno et al. [2014], [32] Sibson and Rowland [2003], [33] Bosl and Nur [2002], [34] Miller et al. [2004], [35] Antonioli et al. [2005], [36] Terakawa et al. [2010], [37] Malagnini et al. [2012], [38] Pezzo et al. [2018], [39] Leclère and Fabbri [2013], [40] Waldhauser et al. [2012].

MCT: Main Central Thrust. Q<sub>P</sub>/Q<sub>S</sub>: attenuation ratios of P and S waves. 3-D: three-dimensional. LVZ: low seismic velocity zone. Φ: stress tensor shape; Φ = (σ<sub>2</sub> − σ<sub>3</sub>)/(σ<sub>1</sub> − σ<sub>3</sub>).
planes [e.g., Sibson, 2013]. The analysis of Seno [2009], based on the construction of the differential stress profiles across several subduction zones, shows that the pore fluid pressure ratio $\lambda$ accounts for much of the interplate coupling.

## 4.2. Aftershocks

### 4.2.1. Present-day aftershocks

In many cases, aftershocks following large magnitude events typically migrate both in space and time. This was first noticed by Nur and Booker [1972] following the fortuitous or induced seismic sequences at Denver and Rangely [Healy et al., 1968, Raleigh et al., 1976]. Since then, precise hypocenter relocations show that the temporal and spatial migration follows laws of diffusion of pressure fronts in porous media, suggesting that pressurized fluids contribute to trigger aftershocks (Table 5). The theoretical analysis of Miller [2020], based on the modeling of permeability fluctuations with time during the seismic cycle, further shows that major events not characterized by co-seismic or post-seismic releases of large amounts of fluids are followed by very few aftershocks, whereas those characterized by the release of significant amounts of co- or post-seismic fluids will be followed by large aftershock populations. Even if other mechanisms such as loading of faults by Coulomb stress transfer can influence the triggering or the propagation of aftershocks [Nostro et al., 2005], the role of supra-hydrostatic $p_f$ appears significant in many cases.

### 4.2.2. Aftershocks in the geological record

Because of their small magnitudes, aftershocks related to past earthquakes are uneasy to recognize in the geological record. However, a specific type of aftershocks, called “golden” aftershocks, is presumably recognized in Archean rocks (Yilgarn craton, western Australia) where mechanical analyses based on Coulomb stress transfer following major earthquakes suggest that transiently abnormally high $p_f$ triggered ruptures on secondary faults after main events [Cox and Ruming, 2004, Micklethwaite, 2008]. Indeed, the Yilgarn craton gold deposits hosted in secondary low-displacement hectometric faults and shear zones do not show a random spatial distribution, but are clustered in rock volumes

near terminations of primary plurikilometric strike-slip fault segments. Mechanical analyses of the stress transfer following major earthquakes along any of the fault segments show that specific rock volumes are stress-loaded ahead of the segment terminations (so-called “loaded” lobes, inside which the stress state acting on pre-existing fractures makes them close to failure; King et al., 1994). Pressurized fluid expelled from the main shock hypocentral volume then percolates through the fracture network in the loaded lobes, triggering eventual ruptures and allowing, by pressure drop, deposition of gold-bearing mineralization. The repetition of main shocks and aftershocks eventually results in deposits of economical interest.

## 4.3. Subducting slab intermediate-depth earthquakes

### 4.3.1. Seismological evidence for supra-hydrostatic $p_f$ in currently subducting slabs

In subduction zones, intermediate-depth earthquakes, that is, earthquakes with hypocentral depths between ~50 and 300 km, are no longer localized along the plate interface, but inside the subducting slab. Various mechanisms have been proposed to account for nucleation or propagation of this intermediate-depth seismicity [Frohlich, 2006, Houston, 2015]: (1) dehydration embrittlement, in which locally abnormally high $p_f$ allows a transition from a viscous regime of deformation to brittle faulting, (2) thermal instabilities, or (self-localizing) thermal runaway, in which incipient ductile shear zones develop as rheological instabilities along which brittle failure can nucleate; (3) “anticrack faulting”, triggered by mineral phase change reactions. Among these mechanisms, which can act in concert, the first one is of peculiar interest here. As summarized in Table 6, evidence or suspicion for supra-hydrostatic $p_f$ come from three-dimensional seismic tomography across subducting slabs currently in their post-seismic stage or in their inter-seismic stage.

### 4.3.2. Evidence for excess $p_f$ at intermediate depths in the geological record

Table 6 further provides examples of structures (crack-seal veins, mineral overgrowths) reported from high-pressure metamorphic terranes which
<table><caption>**Table 6.** Examples of geological and geophysical evidence for intermediate-depth seismicity in modern or fossil subduction zones triggered or favored by supra-hydrostatic pore fluid pressures *p*<sub>f</sub></caption><thead><tr><th>Setting</th><th>Area</th><th>Name of structure</th><th>Evidence</th></tr></thead><tbody><tr><td>NE Japan</td><td>Japan trench (subducting Pacific slab)</td><td>Seismic tomography indicates abnormal V<sub>p</sub>/V<sub>s</sub> and Poisson ratio values interpreted as caused by near-lithostatic p<sub>f</sub><sup>[1]</sup>; seismic tomography shows slab dehydration; released fluids become overpressured and account for seismicity<sup>[2]</sup>; 3D seismic tomography suggests evidence for near-lithostatic p<sub>f</sub> in the hypocentre of a M<sub>w</sub> 7.1 intraslab earthquake at a 66 km depth<sup>[3]</sup>; V<sub>p</sub> and V<sub>s</sub> variations in the subducting slab indicates development of high p<sub>f</sub><sup>[4];[5]</sup></td></tr><tr><td>Central Andean subduction</td><td>Chile trench (subducting Nazca plate)</td><td>High p<sub>f</sub> could have contributed to the nucleation or to the propagation of a M<sub>w</sub> 7.8 earthquake in northern Chile at a 115 km depth<sup>[6]</sup>; high V<sub>p</sub>/V<sub>s</sub> ratios in the subducting Nazca plate and thermodynamic modeling of dehydration reactions suggest high p<sub>f</sub> at depths &gt;50 km<sup>[7]</sup></td></tr><tr><td>Lesser Antilles</td><td>Subducting Atlantic plate</td><td>Hypocenters between 120 and 160 km are distributed where seismic tomography indicates low V<sub>p</sub> values and high V<sub>p</sub>/V<sub>s</sub> ratios, suggesting high p<sub>f</sub><sup>[8]</sup></td></tr><tr><td>Middle America trench</td><td>Middle America trench (subducting Cocos plate)</td><td>Nucleation of a M<sub>w</sub> 7.1 earthquake at a 57 km depth in the subducting Cocos plate tentatively explained by high p<sub>f</sub> resulting from the brucite + antigorite = olivine + H<sub>2</sub>O reaction<sup>[9]</sup></td></tr></tbody></table>

(continued on next page)
<table><caption>**Table 6.** (continued)</caption><thead><tr><th>Setting</th><th>Area</th><th>Name of structure</th><th>Evidence</th></tr></thead><tbody><tr><td rowspan="2">Fossil subduction zones</td><td>Central Alps</td><td>Dent Blanche basal thrust</td><td>Cross-cutting relationships between foliated cataclasites, mylonites and quartz-clinozoisite veins (0.95 &lt; λ &lt; 1)<sup>[10]</sup></td></tr><tr><td></td><td>Valpelline Shear Zone between Valpelline and Arolla units, Dent Blanche</td><td>Metamorphic tensile veins (λ ~ 0.7)<sup>[11]</sup></td></tr><tr><td rowspan="2">Western Alps</td><td>Monviso</td><td>HP crack-seal veins in eclogite-facies metagabbros and oscillatory growth zoning of minerals as products of precipitation suggest fluctuations of p<sub>f</sub><sup>[12]</sup>;</td></tr><tr><td>Lanzo</td><td>eclogite-facies vugs suggest lithostatic p<sub>f</sub><sup>[13]</sup>; four distinct large negative δ<sup>7</sup>Li excursions in garnet mantles result from kinetic fractionation of Li isotopes through bulk diffusion during at least four overpressured fluid pulses<sup>[14]</sup></td></tr><tr><td>California</td><td>Franciscan Complex</td><td>Hydrofracturing of omphacitites by CH<sub>4</sub>-H<sub>2</sub>-rich fluids at 30–80 km depth<sup>[15]</sup></td></tr><tr><td>Zambia</td><td>Relics of subducted lower oceanic crust, late Precambrian Zambezi belt</td><td>Rhythmic chemical zoning in HP garnets result from growth-dissolution cycles driven by pressure pulses (p<sub>f</sub> = 100–350 MPa)<sup>[16]</sup></td></tr></tbody></table>

References: [1] Nakajima et al. [2001], [2] Mishra and Zhao [2004], [3] Nakajima et al. [2011], [4] Shiina et al. [2013], [5] Shiina et al. [2017], [6] Kuge et al. [2010], [7] Bloch et al. [2018], [8] Paulatto et al. [2017], [9] Gutiérrez-Aguilar et al. [2022], [10] Angiboust et al. [2015], [11] Menant et al. [2018], [12] Spandler et al. [2011], [13] Angiboust and Raimondo [2022], [14] Hoover et al. [2022], [15] Giuntoli et al. [2024], [16] Viete et al. [2018], [17] John and Schenk [2006].

HP: high pressure metamorphic conditions.
can be interpreted as resulting from overpressured fluids during intermediate-depth seismic activity. Brittle dilatant veins crossing ductilely deformed high-pressure rocks are one convincing piece of evidence for abnormally high $p_f$. Similarly, finely zoned rims around garnets in eclogitic rocks can be interpreted as repeated growths of thin garnet layers following cyclical arrivals of co-seismic fluids. In these cases, the sharpness of the boundaries between the domains with distinct element concentrations are sharp, whereas they should be irregular or diffuse in the case of intracystalline high-temperature solid diffusion, confirming the fluid pulse hypothesis. Through a detailed analysis of such zoned garnet overgrowths, Viete et al. [2018] estimate that at least four overpressured fluid pulses occurred in less than 300,000 years.

## 4.4. Seismic swarms

Whatever they are related to volcanic systems or not, many seismic swarms appear to be triggered or favored by pressurized fluids (Table 7). The most convincing argument for the role of $p_f$ lies in the fact that, like for aftershock sequences, the spatial and temporal migration of hypocentres follows diffusion laws of pressure front propagation in porous media. Several authors however propose that other mechanisms can also play a role, in combination with $p_f$ wave propagation. A frequently invoked additional mechanism is Coulomb stress transfer [Aoyama et al., 2002, Hainzl, 2004, Yukutake et al., 2011, Fischer et al., 2014]. In particular, studying a swarm developed near the Hakone volcano, central Japan, Yukutake et al. [2011] were able to show that the hypocenters of a first earthquake sequence were aligned along fault-like planar surfaces, while those of a subsequent sequence were located in compressive lobes associated with the planar surfaces, demonstrating the sequence fluid-triggered events followed by Coulomb stress transfer-triggered events. Besides, Aoyama et al. [2002] call for stress corrosion as an additional mechanism in the triggering of seismic swarms.

## 4.5. Fluid overpressures and slow earthquakes

Slow earthquakes, a collective name including slow-slip events, low- or very-low-frequency earthquakes,

low-frequency tremors, non-volcanic tremors, and episodic tremor-and-slip events, are in many cases suspected to be triggered or favored by supra-hydrostatic $p_f$. Table 8 is a non-exhaustive compilation of studies that relate slow earthquakes (in the broad sense) with evidence for supra-hydrostatic $p_f$. With the exception of one study pertaining to the San Andreas fault, all other studies deal with slow earthquakes in subduction zones. Direct evidence come from measurements in ODP or IODP boreholes equipped with pressure or fluid flow sensors [Brown et al., 2005, Araki et al., 2017]. Indirect evidence are provided by geophysical methods (seismic tomography or magneto-tellurics) that show that domains characterized by seismic velocity anomalies or by low resistivities, which are explained by the presence of possibly overpressured fluids, overlap zones of nucleation of slow events. Besides, laboratory experiments carried out on oceanic crust rocks (gabbros, serpentinites) and numerical modeling further support the interplay between supra-hydrostatic $p_f$ and slow earthquakes [Peacock, 2009, Katayama et al., 2012, Kitajima and Saffer, 2012, Beeler et al., 2013, Saffer and Wallace, 2015, Bernaudin and Gueydan, 2018, Condit et al., 2020, Dal Zilio and Gerya, 2022, Eberhard et al., 2022].

### 4.5.1. Evidence in the geological record

Tracking slow earthquake phenomena in the geological record is not easy, since rocks only exhibit frozen structures, for which strain rates are difficult to estimate. However, several studies contend that past slow slip events left tracks in surface-exhumed fault rocks (Table 8). Most of such studies pertain to ancient subduction interface-related slow events triggered or facilitated by near-lithostatic $p_f$. These studies are based on scenarii that are compatible with cyclic fracturing and fluid arrival in the fractured zones. One key issue in the slow-earthquake interpretation is the estimated duration of events. Based on silica diffusion from wall rock to quartz crak-seal veins found along a shear zone adjacent to a paleo-décollement in the Kodiak accretionary complex, Fisher and Brantley [2014] could estimate the time necessary to precipitate quartz for one crak-seal band as less than 10 days, which could correspond to the duration of some slow events in subduction zones. Similarly, based on a kinematic model of the growth of microscopic quartz bands following
<table><caption>**Table 7.** Examples of geophysical evidence for seismic swarms or seismic swarm-like sequences triggered or favored by abnormally high pore fluid pressures *p*<sub>f</sub> in or around fault zones</caption><thead><tr><th>Area and/or events</th><th>Evidence</th></tr></thead><tbody><tr><td>Central and northern Japan (Matsushiro, Tohoku, Hida mountains, Hakone volcano)</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[1],[2],[3]</sup>; pressurized fluid diffusion accounts for space-time evolution of seismicity along with Coulomb stress transfert and stress corrosion<sup>[4]</sup>; pressurized fluid diffusion and Coulomb stress transfert account for space-time evolution of seismicity<sup>[5]</sup></td></tr><tr><td>Dobi graben seismic sequence, Afar rift, Djibouti</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[6]</sup></td></tr><tr><td>East African Rift, Uganda, Rwenzori region</td><td>Spatial overlap in the middle crust between intense volatile and CO<sub>2</sub> circulation and swarm hypocentral regions<sup>[7]</sup></td></tr><tr><td>Central Europe (West Bohemia/Vogtland),</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[8],[9]</sup>, but static and dynamic Coulomb stress changes also play a role<sup>[10],[11],[12]</sup>; at about 9 km depth, λ ~ 0.98 (p<sub>f</sub> = 244 MPa, that is 5 MPa below σ<sub>v</sub>)<sup>[13]</sup></td></tr><tr><td>Western Europe (Vosges mountains)</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[14]</sup></td></tr><tr><td>Western Alps (Ubaye region and vicinity)</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[15],[16],[17]</sup>; supra-hydrostatic p<sub>f</sub> are necessary for reactivation of misoriented faults (λ = 0.41–0.51)<sup>[18]</sup>; time-space evolution of seismicity and development of excess p<sub>f</sub> (between 35 and 55 MPa) explained by creep compaction<sup>[19]</sup>; space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws, but co-seismic stress transfer explains the seismicity close to the mainshock source<sup>[20]</sup></td></tr><tr><td>Corinth rift</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[21],[22]</sup></td></tr><tr><td>Southern Italy, Mt Pollino region</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[23]</sup></td></tr><tr><td>Southern California (including Long Valley caldera)</td><td>Space-time evolution of seismicity fits with p<sub>f</sub> diffusion laws<sup>[24],[25]</sup>, possibly combined with episodic aseismic slip<sup>[26]</sup></td></tr></tbody></table>

References: <sup>[1]</sup>Nur [1974], <sup>[2]</sup>Cappa et al. [2009], <sup>[3]</sup>Yoshida et al. [2016], <sup>[4]</sup>Aoyama et al. [2002], <sup>[5]</sup>Yukutake et al. [2011], <sup>[6]</sup>Noir et al. [1997], <sup>[7]</sup>Lindenfeld et al. [2012], <sup>[8]</sup>Parotidis et al. [2003], <sup>[9]</sup>Hainzl [2004], <sup>[10]</sup>Hainzl and Fischer [2002], <sup>[11]</sup>Fischer and Horálek [2005], <sup>[12]</sup>Fischer et al. [2014], <sup>[13]</sup>Vavryčuk [2002], <sup>[14]</sup>Audin et al. [2002], <sup>[15]</sup>Jenatton et al. [2007], <sup>[16]</sup>Godano et al. [2013], <sup>[17]</sup>Baques et al. [2021], <sup>[18]</sup>Leclère et al. [2012], <sup>[19]</sup>Leclère et al. [2013], <sup>[20]</sup>De Barros et al. [2019], <sup>[21]</sup>Duverger et al. [2015], <sup>[22]</sup>De Barros et al. [2020], <sup>[23]</sup>De Matteis et al. [2021], <sup>[24]</sup>Prejean et al. [2003], <sup>[25]</sup>Shelly et al. [2016], <sup>[26]</sup>Vidale and Shearer [2006].

cyclical pore fluid ingression in crack-seal veins from the Shimanto accretionary prism, Ujiie et al. [2018] could estimate the minimum duration of precipitation between two successive fluid pulses as less than 5 years. Such a duration is too low for standard earthquake rupture, but could correspond

to episodic tremor and slip. Overall, the recognition of slow earthquakes in the geological record is still speculative, but quantification of duration of various inter-seismic or co-seismic phenomena such as mineral precipitation could validate the inferred geological signature of slow events.
<table><caption>**Table 8.** Examples of geological or geophysical evidence for present-day or past slow earthquakes triggered or favored by abnormally high pore fluid pressures *p*<sub>f</sub> in or around fault zones</caption><thead><tr><th>Setting</th><th>Area, structure, unit</th><th>Evidence</th></tr></thead><tbody><tr><td rowspan="3">Present-day events along plate interfaces, subduction zones</td><td>Cascadia</td><td>Seismic imaging and seismic wave analyses suggest that p<sub>f</sub> is near-lithostatic in a LVL inside which SSEs or ETSs nucleate<sup>[1,12], [3], [4]</sup>; supra-hydrostatic p<sub>f</sub> inferred in anisotropically permeable foliation parallel to the plate interface explains NVT distribution<sup>[5]</sup>; 3-D seismic tomography and magneto-tellurics show that LFEs nucleate in a landward dipping region characterized by a high Poisson ratio and by a high electrical conductivity; the region is interpreted as saturated by fluids trapped at near-lithostatic p<sub>f</sub><sup>[6]</sup>; variations in seismic velocities in the LVL after ETSs likely reflect p<sub>f</sub> fluctuations<sup>[7]</sup></td></tr><tr><td>Middle America trench</td><td>Episodic fluid flow pulses measured in ODP boreholes correlate with seismic tremor in the frontal part of the plate interface and could reflect expulsion of transiently overpressured fluids<sup>[8]</sup>; SSEs nucleate in an ultra-slow velocity layer inside which fluids may be overpressured<sup>[9]</sup>; LFE occurrence during a SSE is interpreted to result from a p<sub>f</sub> fluctuation having migrated updip along the subduction interface<sup>[10]</sup></td></tr><tr><td>NE Japan (Japan trench)</td><td>In the upper part of the downgoing Pacific, between 60 and 80 km depth, a zone of high V<sub>p</sub>/V<sub>s</sub> is accounted for by high p<sub>f</sub><sup>[11]</sup></td></tr><tr><td></td><td>SW Japan (Nankai subduction zone)</td><td>NVDT possibly linked with high p<sub>f</sub><sup>[12]</sup>; LFT suppressed where p<sub>f</sub> is low<sup>[13]</sup>; slow slip is located in a high reflectivity and high Poisson ratio portion of the subducting oceanic crust<sup>[14]</sup>; NVT are located in a portion of the subducting oceanic crust characterized by V<sub>p</sub> and V<sub>s</sub> anomalies related to high p<sub>f</sub><sup>[15], [16]</sup>; SSE occur in a portion of the subducting slab characterized by seismic velocity anomalies likely caused by high p<sub>f</sub><sup>[17]</sup>; slow seismic slip in the shallow parts of the subduction interface is characterized by waves rich in high-frequency components, suggesting p<sub>f</sub>-controlled tensile fractures<sup>[18]</sup>; p<sub>f</sub> fluctuations measured in IODP boreholes correlate with SSEs in the shallow parts of the subduction interface<sup>[19], [20]</sup></td></tr></tbody></table>

(continued on next page)
<table><caption>**Table 8. (continued)**</caption><thead><tr><th>Setting</th><th>Area, structure, unit</th><th>Evidence</th></tr></thead><tbody><tr><td></td><td>New Zealand<br>(Hikurangi margin)</td><td>Deep SSEs occur in a zone characterized by a high P wave attenuation and by a high V<sub>P</sub>/V<sub>S</sub> ratio, suggesting high p<sub>f</sub><sup>[21]</sup>; anomalously-shallow slow slip events occur in a low V<sub>P</sub> zone where p<sub>f</sub> is supposed to be supra-hydrostatic (λ* = 0.87)<sup>[22]</sup>; 3-D V<sub>P</sub> and V<sub>P</sub>/V<sub>S</sub> tomography from 8 to 70 km depth along the subduction interface, where both large shallow SSEs and small deep SSEs were recorded, is interpreted as subducted sediment with high p<sub>f</sub><sup>[23]</sup>; OBS data show that tectonic stresses in the subducting slab change throughout SSE cycles, possibly as a consequence of p<sub>f</sub> fluctuations in the vicinity of the plate interface<sup>[24], [25]</sup>; repeating SSEs occur in a zone below which high V<sub>P</sub>/V<sub>S</sub> values suggest high p<sub>f</sub><sup>[26]</sup></td></tr><tr><td>Present-day events along plate boundary transform faults</td><td>San Andreas fault</td><td>NVT activity near Parfield following large earthquakes seem to be enhanced by static shear stress and Coulomb stress increases, but p<sub>f</sub> fluctuations at depth (15–30 km) may also play a role<sup>[27]</sup>; NVT near Parkfield likely reflect shear failure on a critically stressed fault in the presence of near lithostatic p<sub>f</sub><sup>[28]</sup></td></tr><tr><td rowspan="4">Evidence in the geological record along subduction zone plate interfaces</td><td>Namibia (Damara belt)</td><td>Hydrothermal quartz veins preserved in ductilely deformed rocks attest for hydrofracturing caused by high p<sub>f</sub> during LFEs and VLFEs in the plastic aseismic regime below the seismogenic zone<sup>[29]</sup></td></tr><tr><td>Cyclades (Cycladic blueschist unit, Syros)</td><td>Brittle deformation structures (shear fractures and veins) in eclogite lenses embedded in ductilely deformed blueschist are interpreted to result from SSEs triggered by high p<sub>f</sub><sup>[30]</sup></td></tr><tr><td>SW Japan (Triassic Tomuru blueschists)</td><td>Quartz-filled crack-seal shear and extension veins in subduction mélange formed by frictional sliding and tensile fracturing at near-lithostatic p<sub>f</sub> at short recurrence times (a few years, determined from kinetic modeling of quartz precipitation) are tentatively accounted for by episodic strain release during ETSs<sup>[31]</sup></td></tr><tr><td>New Zealand (Dun ophiolite, Livingstone Fault)</td><td>Field and microstructural observations suggest that chemical reactions involving serpentinite can promote rock hardening and generate in situ fluid over-pressure patches and brittle failure in the source region of deep tremor along the slab-mantle interface<sup>[32]</sup></td></tr></tbody></table>

(continued on next page)
<table><caption>**Table 8.** (continued)</caption><thead><tr><th>Setting</th><th>Area, structure, unit</th><th>Evidence</th></tr></thead><tbody><tr><td></td><td>Alaska (Kodiak accretionary complex)</td><td>Quartz vein microstructures and estimation of duration of local silica diffusion indicate that near-lithostatic p<sub>f</sub> can produce crack seal bands in less than 10 days, which is similar to the duration of slow earthquakes<sup>[33]</sup></td></tr><tr><td></td><td>Northern Apennines (continental metasediments)</td><td>Syn-mylonitization quartz-carpholite dilational shear veins formed under HP metamorphic conditions as a result of high p<sub>f</sub> are interpreted as a record of deep ETS events<sup>[34]</sup></td></tr><tr><td></td><td>Central Alps (Arosa zone)</td><td>Crack-seal veins formed under blueschist facies conditions concurrently with viscous flow as a result of near-lithostatic p<sub>f</sub>; the low estimated values of differential stresses suggest that the veins formed during slow slip and tremor events<sup>[35]</sup></td></tr><tr><td></td><td>California (Franciscan complex eastern belt)</td><td>Opening increments of ~1 mm in blueschist-facies metasediments are interpreted as the result of transient near-lithostatic p<sub>f</sub> during LFEs<sup>[36]</sup></td></tr></tbody></table>

References: [1] Abers et al. [2009], [2] Audet et al. [2009], [3] Tauzin et al. [2017], [4] Audet and Schaeffer [2018], [5] Delph et al. [2018], [6] Calvert et al. [2020], [7] Gosselin et al. [2020], [8] Brown et al. [2005], [9] Song et al. [2009], [10] Frank et al. [2015], [11] Shelly et al. [2006b], [12] Obara [2002], [13] Seno and Yamasaki [2003], [14] Kodaira et al. [2004], [15] Shelly et al. [2006a], [16] Matsubara et al. [2009], [17] Kato et al. [2010], [18] Sugioka et al. [2012], [19] Araki et al. [2017], [20] Ariyoshi et al. [2021], [21] Wallace and Eberhart-Phillips [2013], [22] Bassett et al. [2014], [23] Eberhart-Phillips and Bannister [2015], [24] Shaddox and Schwartz [2019], [25] Warren-Smith et al. [2019], [26] Yarce et al. [2021], [27] Nadeau and Guilhem [2009], [28] Thomas et al. [2009], [29] Fagereng et al. [2018], [30] Behr et al. [2018], [31] Ujiie et al. [2018], [32] Tarling et al. [2019], [33] Fisher and Brantley [2014], [34] Giuntoli and Viola [2021], [35] Condit and French [2022], [36] Schmidt and Platt [2022].

IVL: low-velocity layer. SSE: slow-slip event. ETS: episodic tremor and slip. NVT: non-volcanic tremor. 3-D: three-dimensional. LFE: low-frequency earthquake. NVDT: non-volcanic deep tremor. ODP: Ocean Drilling Program. LFT: low-frequency tremor. OBS: ocean-bottom seismometer. VLFE: very low frequency earthquake. HP: high pressure.
## 4.6. Fluids and man-made earthquakes

Several man-made earthquakes show that fluids can trigger seismic ruptures. One emblematic piece of evidence for this triggering role is the earthquake activity that occurred in Colorado following waste water disposal at depth [Healy et al., 1968]. Experimental water injection at depth at Rangely [Colorado, Raleigh et al., 1976] confirmed than fluids could trigger earthquakes. Since then, the artificial triggering of earthquakes has been ascertained in a large variety of human activities: hydrocarbon extraction [Wiprut and Zoback, 2000, Amos et al., 2014, Bourne et al., 2014, Van Wees et al., 2014], dam reservoir impoundment [Rastogi et al., 1997, Gupta, 2002, McGarr et al., 2002, Zhang et al., 2019], industrial waste water repository [Horton, 2012, Yeck et al., 2016], carbon dioxide storage projects [Zoback and Gorelick, 2012], and water injection in deep boreholes to produce geothermal energy [Deichmann and Giardini, 2009, Terakawa et al., 2012, McGarr et al., 2015]. Fluid-injection-experiments at the KTB borehole in Germany have shown that even small pressure variations (<1 MPa) could trigger numerous microearthquakes at a depth of 9 km [Zoback and Harjes, 1997]. In most case studies, the $p_f$ build-up in the vicinity of critically loaded faults is responsible for failure. However, abnormally high $p_f$ may also act indirectly by triggering aseismic slip which in turn leads to an earthquake [Wei et al., 2015]. Lastly, stress perturbations by Coulomb stress transfer can also contribute to man-induced seismic ruptures [De Matteis et al., 2024].

## 4.7. Conclusions

The interactions between fluids and seismic ruptures in the brittle crust or in subducting slabs are more and more recognized or suspected (Tables 5–8). The role of overpressured fluids in the triggering or propagation of small-magnitude events such as aftershocks, swarms and, to a lesser extent, slow earthquakes, seems to be well established. Regarding large-magnitude events, the role of excess pore fluid pressures is more difficult to ascertain. However, a large body of evidence suggests that seismic ruptures, especially those along plate interfaces in subduction zones, take place in crustal volumes where seismic tomography indicates body wave velocity anomalies that are interpreted as the consequence of

the presence of overpressured fluids. The present review of literature further shows that abnormally high pore fluid pressures may not be the only mechanism. Indeed, stress perturbations can act in concert with abnormally high $p_f$ and the inferred decrease of $\sigma_n$.

## 5. Conclusions

A review of the literature shows that fluids physically interact everywhere and at all scales with deformation in the brittle crust and in subducting slabs. Since they are more permeable than most intact rocks, fractures, faults and shear zones in the brittle or brittle-ductile crust and in the subducting slabs appear as efficient paths for fluid circulation. Provided that impervious seals develop transiently or permanently, fluids can accumulate and their pressure can increase. Where faults are critically stressed, only a minor overpressure value (a few MPa above the hydrostatic pressure) can trigger or favor rupture or displacement. Where faults are not critically stressed, for instance where they are severely misoriented with respect to the active stress tensor, larger excess $p_f$ values are needed for reactivation. Fluid overpressures can be estimated by using chemical data or fluid inclusion data in exhumed rocks. Estimated values are as follows: 25–135 MPa at seismogenic depths [ca. 5 to 10 km, Shimanto paleo-accretionary prism, Raimbourg et al., 2022, Kodiak paleo-accretionary prism, Vrolijk, 1987], 50–150 MPa at 8–9 km [Apenines, Mullis, 1988], 110 MPa at ca. 20 km [Central Alps, Berger and Herwegh, 2019], 200–350 MPa at 8–10 km or deeper [Val d'Or district, Canada, Robert et al., 1995], 100–350 MPa at 30–70 km in an exhumed subducted lithosphere [Viete et al., 2018].

The influence of abnormally high $p_f$ is of utmost importance in the triggering or in the propagation of seismic ruptures. The $p_f$ effect appears to be common in small to moderate magnitude ($M < 5 \sim 6$) events such as aftershocks or swarm sequences. The role of high $p_f$ is less obvious for large magnitude ($M > 6$ or higher) events. However, the precursor mechanisms preceding large magnitude events may be partly controlled by $p_f$ build-up. Future researches will concentrate on this aspect of the seismic cycle. In particular, a critical issue is to try to determine whether a $p_f$ increase is the triggering (causal) mechanism for a seismic rupture propagation or is a consequence of a propagating
rupture triggered by a mechanism that would not be related to $p_f$. In their spatio-temporal monitoring of $V_p$ and $V_s$ variations during the 1997 Umbria-Marche sequence, central Italy (Table 5), Chiarabba et al. [2009a] suggest that the $p_f$ propagation precedes the seismic ruptures. A similar conclusion is reached by Yoshida et al. [2023] in their study of a swarm sequence resulting in a $M_w$ 6.2 event in NE Japan. Besides, the *in-situ* experiments of Guglielmi et al. [2015] and Cappa et al. [2022] show that a $p_f$ increase triggers aseismic slip that eventually turns into a seismic displacement. From their field study of intermediate-depth earthquakes frozen in an ancient oceanic crust, John and Schenk [2006] conclude that seismic rupture was not preceded but rather accompanied and followed by fluid ingression in the hypocentral volume (Table 6). Further experimental or *in situ* works along with detailed geological or geophysical analyses are obviously needed before the complex interrelationships between $p_f$ variation and inception or propagation of brittle fracture or slip along faults can be fully understood.

High $p_f$ values also play a role in the triggering or maintenance of slow slip earthquakes and related phenomena along several subduction interfaces or along transform faults. Since these transient phenomena can in turn influence or control large-magnitude seismic ruptures, a careful analysis or monitoring of $p_f$ build-up is an important direction of future research. The precursory results of Brown et al. [2005], Mikada et al. [2006], Fulton and Brodsky [2016] or Ariyoshi et al. [2021] are examples of such a promising monitoring.

Declaration of interests

The authors do not work for, advise, own shares in, or receive funds from any organization that could benefit from this article, and have declared no affiliations other than their research organizations.

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