# Swarms and mainshock-aftershocks sequences are both triggered by fluids in the Ubaye Region (Western Alps)

Marion Baques, Louis de Barros, Maxime Godano, Clara Duverger, Hervé Jomard

► To cite this version:

Marion Baques, Louis de Barros, Maxime Godano, Clara Duverger, Hervé Jomard. Swarms and mainshock-aftershocks sequences are both triggered by fluids in the Ubaye Region (Western Alps). Geophysical Journal International, 2023, 235 (1), pp.920-941. 10.1093/gji/ggad280. hal-04182990

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Swarms and mainshock-aftershocks sequences are both triggered

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by fluids in the Ubaye Region (Western Alps).

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Marion Baques¹, Louis De Barros¹, Maxime Godano¹, Clara Duverger² and Hervé Jomard³

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1: Université Côte d'Azur, CNRS, Observatoire de la Côte d'Azur, IRD, Géoazur, 250 rue Albert

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Einstein, Sophia Antipolis 06560 Valbonne, France

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2: CEA, DAM, DIF, F-91297 Arpajon, France

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3: Bureau d'évaluation des risques sismiques pour la sûreté des installations, Institut de

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Radioprotection et Sûreté Nucléaire, Fontenay-aux-Roses, France

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**Abbreviated title:**

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Fluid-triggered mainshock-aftershocks and swarms

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**Corresponding author:** Baques Marion Mail: [marion.baques@geoazur.unice.fr](mailto:marion.baques@geoazur.unice.fr)

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21 Abstract

The Ubaye Region (French Western Alps) is one of the most seismically active regions in France. It is regularly struck by mainshock-aftershocks sequences like in 1959 (ML5.3), seismic swarms (2003-2004), and complex sequences (2012-2015) characterized by successive mainshocks clustered in time and space. This diversity of seismic behaviour highlights the complex processes at play in this area.

To improve our understanding of these processes, in this study, we compile a regional catalogue of existing focal mechanisms, completed by 100 new calculated focal mechanisms of aftershocks following the 07/04/2014 mainshock (ML5.1). The oriented stress-state we reconstruct for different periods and sub-areas are similar to each other and to previous published values focusing on swarm periods. We then calculate fluid-pressure required to trigger the earthquakes. Most of the events (65%) need fluid-overpressure between 15 and 40 MPa (17-to-40% of the hydrostatic pressure) with a median value of 24%. Moreover, even the largest events, like the mainshocks in the 2012-2015 sequence, appear to be triggered by fluid-pressure, similarly as events within swarm sequences.

However, while fluid-overpressure decreases with time in an aftershock sequence, it varies randomly at high levels during a swarm sequence. Therefore, based on a fault-valve model, we propose that: 1) the fluids trapped in the fault plane tend toward lithostatic pressure and trigger the mainshock rupture and 2) part of the aftershocks are induced by the diffusing fluid-pressure. On the contrary, swarms need external, likely deep, fluid-pressure feedings. Fluid-pressure is likely to be a common triggering mechanism of the seismicity in the Ubaye Region, even if the involved processes should differ to explain the different types of seismic sequences.

41 Keywords

Europe; Earthquake dynamics; Seismicity and tectonics; Dynamics and mechanisms of faulting; Fluid-driven seismicity

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# 1. Introduction

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Two end-member behaviours may characterize earthquake sequences: mainshock-aftershocks sequences and swarm sequences. On one hand, mainshock-aftershocks sequence is the most common behaviour and starts with the largest magnitude earthquake, possibly damaging, followed by aftershocks of lower magnitude. Aftershocks are mainly driven by coseismic stress transfer (Stein, 1999) and occur around the fault zone where the mainshock occurred. However, some of the aftershocks sequences may also have been triggered by fluid-pressure, reactivating fault by high fluid-pressure (e.g. Gunatilake, 2023). The number of aftershocks decays with time following an Omori-Utsu's law (Utsu and Ogata, 1995). Such mainshock-aftershocks sequences occur in all tectonically active contexts, as the tectonic stress loading directly triggers them. On the other hand, swarm sequences are bursts of earthquakes, usually of small-to-moderate magnitudes, which are clustered in time and space. They are characterized by a slow increase and decrease of the seismicity rate without a clear pattern, and by the strongest event occurring at any time during the sequence. They also occur in different tectonic settings: subduction zones (e.g. Japan, Kato et al. 2014; Nishikawa and Ide, 2017), volcanic regions (e.g. Mayotte, Bertil et al. 2021) or stable intraplate areas (e.g. Western Bohemia, Fischer et al. 2014). They can last for days (e.g. Mammoth Mountain, Shelly and Hill, 2011; Mount Rainier, Shelly et al. 2013; Corinth Gulf, De Barros et al. 2020; South Korea, Son et al. 2021) to years (e.g. Ubaye, Jenatton et al. 2007; Cahuilla swarm, Ross and Cochran, 2021; Japan, Amezawa et al. 2021). To maintain this activity, a driving mechanism that lasts in time is required. As swarms can also be induced by geothermal activities (e.g. Baisch et al., 2015; Schoenball et al., 2014) or hydrocarbon resources exploitations (e.g. Goebel et al. 2016; Schultz et al. 2020) or even triggered by rainfalls during monsoon period (e.g. Gunatilake et al., 2021), they are commonly
associated with fluid-pressure diffusion at depth (e.g. Miller et al. 2004; Parotidis and Shapiro, 2004; Terakawa et al. 2010). Classically, seismic failures are seen to be induced directly by the decrease of normal effective stress associated with the increase of fluid-pressure. Seismic migration is then explained by the fluid-pressure diffusion. Alternatively, swarm sequences were also observed to be driven by slow-slips or aseismic deformation (Lohman and McGuire, 2007; Vallée et al. 2013). Fluid and aseismic processes may also be imbricated to drive swarms as the first one drives the latter, as observed in different contexts (Corinth Gulf, De Barros et al. 2020; Canada, Eyre et al. 2020) or numerical simulations (e.g. Wynants-Morel et al. 2020). As no direct observations are possible in the source area, swarm processes are still an open question.

In between these two end-member categories, more complex sequences exist. For example, some sequences involve several cascading mainshocks, with a global swarm pattern (Colfiorito sequence, Chiaraluce et al. 2003; Ubaye region, Thouvenot et al. 2016; Western Bohemia, Hainzl et al. 2016, Jakoubková et al. 2018; Corinth Gulf, Kaviris et al. 2021). In the Ubaye Region (De Barros et al. 2019) as well as in the Puerto Rico area (Ventura-Valentin and Brudzinski, 2021) or the Corinth Gulf (Kaviris et al. 2021), authors observe that a mainshock started a seismic crisis, but with an abnormal number of aftershocks that did not decay according to the expected Omori-Utsu's law and contained small swarm sequences. In such cases, complex processes involving tectonic stress release as well as fluids and/or aseismic-slip are required. The Amatrice-Visso-Norcia sequence (Appenines, Italy) was also a complex seismic sequence during which the generation of internal fluids by decarbonisation seems to be responsible for non-Omori type aftershock behaviour (Gunatilake and Miller, 2022).

The Ubaye Region, located in the Southwestern Alps, is one of the most seismically active area in France. Mainshock-aftershocks, seismic swarms, and complex sequences have been observed since the 18th century (Baques et al. 2021). In 2003-2004, more than 16,000 events were recorded in a swarm sequence (Jenatton et al. 2007; Daniel et al. 2011; Leclère et al. 2013). The last seismic crisis happened in 2012-2015 with two moderate earthquakes of magnitude ML4.8 and ML5.1, which
occurred at a similar location on 26/02/2012 and 07/04/2014, respectively. They were followed by an abnormally high rate of seismicity (Thouvenot et al. 2016; De Barros et al. 2019), including several small swarms at close distances from the mainshocks. Globally, this seismic crisis can be described as a swarm, because the size of the seismic cluster is much larger than the spatial extent of the main ruptures. As 1) the tectonic deformation is very limited in the Southern Alps (e.g. Mathey et al. 2021b; Walpersdorf et al. 2018), and 2) the areas just South or North of the Ubaye Valley are less active seismically, complex processes are required to explain both the high rate of seismicity and its complex patterns. For both the 2003-2004 swarm and the 2012-2015 sequence, fluid processes have been raised (e.g. Leclère et al. 2013; Thouvenot et al. 2016; De Barros et al. 2019). For example, seismic migration with a diffusivity of 0.002 to 0.06 m²/s were observed for the 2003-2004 and within the 2012-2015 sequences (Thouvenot et al. 2016; De Barros et al. 2019). These values are similar to the measured ones for fluid-induced seismicity in reservoir exploitations (Shapiro et al. 2002) and for other swarms worldwide (Talwani et al. 2007). Leclère et al. (2013) proposed that creep compaction pressurized the fluid within the fault, which then drove the 2003-2004 swarm migration. Daniel et al. (2011) also proposed that fluid process is required, because stress-transfer through earthquake-earthquake interaction explains only 59% of the 2003-2004 swarm. De Barros et al. (2019) suggested an interplay between stress-transfer and fluid-driven process since only some clusters in the 2014 aftershocks followed a typical aftershocks pattern, while others clearly behaved as seismic swarms.

In this study, we aim at better understanding the processes at play, focusing both on the 2014 aftershocks period that was recorded by a dense temporary network (e.g. Baques et al., 2021) and on specific seismic sequences, such as the 2003-2004 seismic swarm and the 2012-2015 seismic crisis, to explain the peculiar seismic behaviour in the Ubaye Region.
Direct observations of fluid are not possible at crustal depth. However, as fluid-pressure induces failures through reduction of the effective stress, it may trigger seismicity on fault planes that are misoriented with respect to the tectonic stress-state (Sibson, 1990). We therefore aim to compare the orientation of the observed failures, obtained through focal mechanisms, with the average stress-state. Using Cauchy's equations, we can reconstruct the fluid-pressure that is required within the faults to induce earthquakes. A similar approach was used by Leclère et al. (2013), but focusing only on the 2003-2004 swarm period. They showed that the seismicity required from 20 to 50 MPa of overpressure with ~35 MPa at the beginning of the swarm and increasing up to ~55 MPa during the peak of activity.

After presenting the general seismo-tectonic context of the area, we first describe the different focal mechanisms catalogues we use for the Ubaye Region, including the well-instrumented two-month period after the 2014 mainshock. We then present the methods used to reconstruct the stress-state and the fluid-pressure needed to trigger the seismicity. We apply these methods on the Ubaye Region and refine the results looking at sub-areas, for different time periods, and for the 2014 aftershocks. Finally, we propose and discuss a conceptual model that explains the role of fluid-driven process in the Ubaye Valley area.

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## 2. Seismo-tectonic context

### 2.1. Tectonic context

The Alpine chain is divided into two zones, internal and external, delimited by the Frontal Penninic Thrust (FPT, Fig. 1). East of the FPT, in the intern zone, the Ubaye Region is mainly composed of the
Briançonnais unit, which consists of a pile of thick low grade metamorphic nappes (e.g. Sue et al. 2007). West of the FPT, the Dauphinois units is composed of a 1-to-2-km thick Mesozoic to Cenozoic sedimentary rock cover (sandstone, marls, and limestone, Kerckhove et al. 1978) laying on the crystalline basement (granitoid; Baietto et al., 2009). Overlaying some parts of the Dauphinois units is the 1-to-2 km thick "Flysch à Helminthoïdes". This cover is also called the Embrunais-Ubaye nappes (Kerckhove, 1969) and is composed of late Cretaceous turbiditic series that were thrust westward during the Oligocene period (e.g. Gratier et al. 1989).

Crossing perpendicularly the Alpine structures, the Ubaye Valley gives us a view of the long geological history of compression and extension resulting from the setting of the Western Alps. The late Alpine faults as defined by Kerckhove (1969) and Tricart (2004) are the more recently recorded deformations. It consists in a series of normal and right lateral strike-slip faults formed or reactivated during the last Mio-Pliocene setting of the chain, and showing field evidence of neotectonic activity (Sue, 1998; Sue et al. 2007). Current seismicity seems to develop along them (Jenatton et al. 2007; Sue et al. 2007; Le Goff et al. 2009; Sanchez et al. 2010; Mathey et al. 2020; Fig. 1). A majority of these faults are oriented N-S to NW-SE with some minor conjugate structures oriented E-W to NE-SW (Kerckhove, 1969; Sue and Tricart, 2003; Tricart and Schwartz, 2006). The main identified active faults in the Ubaye Region are the High-Durance, the Serenne, and the Jausiers-Tinée faults (Fig. 1; Baques et al. 2021; Mathey et al. 2020). The first is oriented N-S and is prolonged to the Southeast by the Serenne fault. These two faults are located east of the Frontal Penninic Thrust while the Jausiers-Tinée fault (external part) is parallel to and at about 10 km south of the Serenne fault.

Using satellite geodesy, there is no measurable horizontal strain (e.g. Walpersdorf et al. 2018). For the vertical deformation, Mathey et al. (2021b) used INSAR data to obtain the vertical strain-rate in the Western Alps and found a range of ~-0.7mm/yr to ~0.5mm/yr for the Ubaye Region (see Fig. 3 in Mathey et al. 2021b). Therefore, vertical deformation is very limited and not homogenous in the Ubaye Region.
Several studies computed earthquake focal mechanisms and calculated stress-state in the Ubaye Region and its surrounding areas to determine the tectonic stress regime. The mechanisms showed a dominant extensional regime with a strike-slip component for the Ubaye Region (Fig. 1). Mathey et al. (2021b) found that the extensional component is oblique to the strike of the Alpine belt, while it was assumed normal until now (Sue et al. 2007). Similar stress-states have been reconstructed for the last swarms (2003-2004 and 2012-2015) occurring in the Ubaye Region (Leclère et al. 2013; Fojtíková and Vavryčuk, 2018).

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

<figure><img src="image_3.png" /><figcaption>Figure 1: a) Simplified geological map of the Ubaye Region with focus on the late Alpine faults. (red lines, from Tricart, 2004). Dots correspond to the location of earthquakes with their local magnitude</figcaption></figure>
179 ML (from the LDG, Duverger et al. (2021)) greater than 3.2, the colour being proportional to ML.

180 Focal mechanisms of the largest events are from several catalogues: Fréchet and Pavoni (1979),

181 Leclère et al. (2013), Thouvenot et al. (2016), FMHex20 (Mazzotti et al. 2021) and Mathey et al.

182 (2021a). Their colours show the depth of the events. The clusters of events between Crévoux,

183 Jausiers, and Barcelonnette (2003-2004 and 2012-2015 swarms) have been artificially dispersed for

184 better visualization. b) Simplified geological cross-section oriented N050°, modified from Dumont et

185 al. (2022). The dots show the events with a computed focal mechanism (Fig. 2b). Black dots are

186 events with magnitude smaller than 3.2, while larger magnitude are coloured according to the scale

187 in a).. Only the events contained in a buffer of 10km apart from the cross-sections are represented.

188 The late Alpine faults are extended at depth by dotted lines for illustration but no information on the
<figure><img src="image_4.png" /><figcaption>Figure 2: a) Map of the instrumental seismicity since 1965 (LDG catalogue, http://www-dase.cea.fr).</figcaption></figure>

The orange and pink dashed-line zones delimit the Ubaye Valley and the Serenne area respectively, and the red square shows the 2014 relocated aftershock area (Fig. 3). b) Map of the earthquakes with a computed focal mechanism in the Ubaye Valley (circle) and the Serenne area (triangle) coloured by the year of occurrence. c) Map of the earthquakes with computed focal mechanisms for the 2003-2004 swarms (light blue circle), the 2012-2015 sequence (dark blue circle), and of the background seismicity (triangle) coloured by the year of occurrence. In b) and c), the red lines
correspond to the late Alpine faults mapped by Tricart (2004): D = High-Durance fault; S = Serenne fault; JT = Jausiers-Tinée fault. d) Distribution of earthquakes in the LDG catalogue as function of the time (years) for two different areas: the Ubaye Valley (orange) and the Serenne area (purple). The dashed black line (Background seismicity) corresponds to events that occurred in the Ubaye Valley during the periods: 1965-2002, 2005-2011, and 2016-2020. The diamonds correspond to the events with magnitude greater than or equal to 3.5.

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## 2.2. Seismic activity

Even if background seismicity is scattered in the full area, most events are located in two main areas: in a large band along the Serenne and High-Durance faults in the internal units later referred as "Serenne area", and in large clusters between Crévoux, St-Paul-sur-Ubaye, and Barcelonnette in the external units later referred as "Ubaye Valley area" (Fig. 2). The seismicity in these two areas behaves differently (Baques et al. 2021; Larroque et al. 2021). In the Ubaye Valley, the seismicity mainly occurs in clusters (Fig. 2a) and shows both mainshock-aftershocks and swarm sequences. On the contrary, in the Serenne area, the seismicity is more scattered without clear clustered zones. The seismicity mainly develops as mainshock-aftershocks sequences. Yet, the presence of small seismic swarms cannot be discarded as no specific studies have been made on this area. In the following, we briefly describe the two clusters that occurred in 2003-2004 and 2012-2015 in the Ubaye Valley area.

The 2003-2004 swarm (Fig. 2, light blue circles) started in early 2003 with seismic activity slowly increasing until July 2003. The seismicity rate remained at a high level for 3 months with the strongest earthquakes (maximum magnitude of ML<sub>LDG</sub>3.4) occurring during the end of the summer 2003. Then the seismicity rate slowly decreased to reach the background seismicity rate in the summer 2004 (Jenatton et al. 2007; Baques et al. 2021). Finally, a small burst of activity occurred at the end of 2004. More than 16,000 earthquakes were detected between 3km and 8km depth

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222 (crystalline basement) and aligned along a 9km long structure oriented N150°E. Leclère et al. (2013)
223 computed 74 focal mechanisms that showed combined right lateral strike-slip motions along N150°-
224 to-N175°E structures with extensional components. Among the focal mechanisms, one of their nodal
225 planes shows similar orientation as the main direction of the cluster (~N150°E, Jenatton et al. 2007)
226 and to the orientation of the Jausiers-Tinée fault system. Leclère et al. (2013) proposed that the
227 Embrunais-Ubaye nappes, as well as the thick Mesozoic marly layers play the role of cap rock,
228 allowing the fluids to pressurize the upper crust and generate earthquake swarms.
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230 Ten years later, the 2012-2015 sequence started 7 km north-west of the 2003-2004 swarm, (Fig. 2,
231 dark blue circles) on 26 February 2012 with a ML<sub>LDG</sub>4.8 (Mw4.2) earthquake at 8.8 km depth. It was
232 followed two years later, on 7 April 2014, by a larger shock of ML<sub>LDG</sub>5.2 (Mw4.8), which occurred
233 exactly at the same location but slightly deeper (9.9 km, Thouvenot et al. 2016). These mainshocks
234 were felt in Nice and Grenoble, 100 km from the epicentre area (Sira et al. 2012, 2014; Courboulex et
235 al. 2013). Following the two mainshocks, intense seismic crises started, and the 2014 crisis
236 instantaneously reactivated the 2012 crisis area. The seismicity then globally decreased as aftershock
237 sequences (Thouvenot et al. 2016; De Barros et al. 2019). During the two months (10 April 2014 - 10
238 June 2014) following the 7 April 2014 mainshock, a temporary network of 7 seismic stations was
239 deployed at 5 to 16 km from the mainshock epicentre. More than 9,000 events were detected by
240 template-matching and ~6,000 of them were relocated (Fig. 3) using a double-difference relocation
241 by De Barros et al. (2019). A modified Omori-Utsu's law (Utsu, 1961) fitted on these events showed a
242 p=0.77 time exponent, which is slower than usual and reflects a more complex behaviour (see Fig. 5
243 in De Barros et al., 2019). This aftershock sequence is interpreted as composed of intertwined
244 aftershock-like clusters and swarm sequences (De Barros et al. 2019). After the 2014 mainshock-
245 aftershocks sequence, an event of ML<sub>LDG</sub>3.6 (Mw=3.4) occurred in mid-April 2015, followed by
246 another of ML<sub>LDG</sub>4.3 (Mw=4.2) in November 2015. Despite some classical mainshock-aftershocks
sequence characteristics, the full crisis (2012-2015) follows a swarm-like behaviour with the largest
event in the middle of the temporal sequence and a spatial extension much larger than the rupture
dimension of the largest event. The majority of the seismicity occurred in the crystalline basement
between 4 km and 10 km depth and aligned along a N160°E major structure and some N50°E minor
structures (Fig. 3, Fig. S1). These orientations are retrieved in the 13 focal mechanisms computed by
Thouvenot et al. (2016). In addition, no seismicity occurred in the area between the 2003-2004
swarm and the 2012-2015 sequence (Thouvenot et al. 2016). The activated structures, well
illuminated at depth (Thouvenot et al. 2016; De Barros et al. 2019) are not identified in the surface
geology, possibly due to the compliant sedimentary materials, which make it difficult to locate and
recognize active deformation. On the contrary, active structures are identified at the surface in the
internal zone (Sue et al. 2007), but the diffuse nature of seismicity does not allow to clearly associate
the seismicity to well identified single fault systems.

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**3. Data: focal mechanisms**

In this study, we based our analysis on existing focal mechanisms computed previously in the Ubaye
Region, and on focal mechanisms specially computed for the two months aftershocks following the 7
April 2014 mainshock.

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**3.1 Ubaye Region**

Thereafter, we gathered all the available focal mechanisms in a catalogue (full Ubaye Region
catalogue). This catalogue is composed of the FMHex20 catalogue from Mazzotti et al. (2021),
mechanisms from Fréchet and Pavoni (1979), Leclère et al. (2013), Thouvenot et al. (2016), and
270 Mathey et al. (2021a). In total, 799 focal mechanisms composed the full Ubaye Region catalogue, 271 those with a magnitude larger than ML<sub>LDG</sub>3.4 (4% of the events) are displayed in Fig. 1. About 39% of 272 the focal mechanisms occurred with a pure strike-slip to strike-slip motion with a normal component, 273 while ~40% of the events show a pure normal to normal with a strike-slip component mechanism 274 (Fig. S2, Tab. S1). The remaining part shows strike-slip with a reverse component (12%) or reverse 275 mechanism (~9%). Two orientations stand out: N145°-to-N250°, N330°-to-N65° (Fig. S3). These 276 orientations are compatible with the active faults in the area. Indeed, the N145°-to-N250° is similar 277 to the orientation of the Serenne fault (N150°), while the N330°-to-N65°direction is consistent with 278 the High-Durance fault and with conjugate faults of the Serenne fault (Fig. 1).

279 We then refine the full catalogue by separating mechanisms from the Ubaye Valley (453 events, Fig. 2b) and from the Serenne area (314 events, Fig. 2b). We also create two subsets of focal mechanisms 280 for the Ubaye Valley: the 2003-2004 swarm area (138 events, Fig. 2c) and the 2012-2015 sequence 281 area (260 events, Fig. 2c). The remaining focal mechanisms in the Ubaye Valley are gathered to form 282 a background seismicity catalogue (55 events, Fig. 2c) that covers the periods of 1965-2002, 2005- 283 2011, and 2016-2020. No significant differences are observed either in the distributions of the fault 284 mechanisms or in the nodal plane orientations among the different sub-catalogues (see Tab.S1 and 285 Fig. S2). 286

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## 3.2 2014 Aftershocks

289 For this sequence, only 3 focal mechanisms were already computed (Thouvenot et al. 2016; Mathey 290 et al. 2021a). We choose to refine the analysis by computing the focal mechanisms of the largest 291 events in the relocated catalogue of De Barros et al., 2019. In total, 15 seismic stations were used: 7 292 from the temporary network and 8 permanent ones (3 from the INGV network (University of Genoa., 293 1967; INGV Seismological Data Centre, 2006); 5 from the RESIF network (RESIF, 1995)). For each
event, we invert simultaneously the P-wave polarity and the S-to-P-waves amplitude ratio to determine the focal mechanisms (Godano (2009), Godano et al. (2014)). For the amplitude ratio, we first picked the amplitude of the first arrival for the P- and S-waves on the three components and calculated the norm for both wave types. We carefully check the consistency of our results with the focal mechanisms obtained by Thouvenot et al. (2016). Examples of focal mechanisms and their uncertainties, together with the P-wave polarities and the S/P amplitude ratio, are shown in Supplementary Fig. S4. We computed the mechanisms for which polarity and amplitude ratio are available from at least 7 stations. At the end of the inversion, we kept solutions for which inversion leads to score (inversed misfit function) greater than 0.4 (see Godano et al. 2014). The 100 focal mechanisms then form the 2014 relocated aftershocks catalogue.

The 2014 relocated aftershocks catalogue is mainly composed of 70% of pure-normal to normal mechanisms with a strike-slip component and 20% of strike-slip mechanisms with a normal component (Tab.S1 and Fig. S2). Even if the ratio of normal mechanisms is higher than in the Ubaye Valley catalogue (Fig. S2), we retrieve similar orientations for the nodal planes (N170°-to-N230°, N355°-to-N70°; Fig. S3). However, we do not retrieve the N155° orientation proposed for the 2014 mainshock. Globally, the 2014 catalogue shows a wide variety of focal mechanisms and orientations, which suggests that complex processes may be in play in this aftershock sequence (Fig. 3). Finally, as the Ubaye catalogue and the 2014 aftershocks come from different origins, they will not be used together, but in separate inversion process, for the stress-state reconstruction and the fluid-pressure computation.
<figure><img src="image_5.png" /><figcaption>Figure 3: Map of the aftershocks relocated by De Barros et al. (2019) in the two months following the 7 April 2014 mainshock, with a subset of 32 (out of 100) computed focal mechanisms. The colours show the event depths. The purple star corresponds to the location of the 2014 mainshock with its focal mechanisms from Thouvenot et al. (2016).</figcaption></figure>
# 4. Methods

## 4.1 Shape and orientation of the stress tensor

To estimate the fluid-overpressure required to trigger earthquakes, we use the focal mechanisms to reconstruct the stress-state orientation. To do that, we use the MSATSI program (Martínez-Garzón et al. 2014), which is a MATLAB version of the SATSI program developed by Hardebeck and Michael (2006). It is based on the Wallace (1951) and Bott (1959) assumption: the slip vector is assumed parallel to the maximal shear-stress. The stress-state is then determined through a grid-search inversion in order to maximize the shear-stress acting on the slip direction for all earthquakes. The MSATSI software requires as input the dip direction (strike+90°), the dip angle, and the rake, and computes the orientation of the principal stresses and the shape-ratio $\varphi$ (Michael, 1984):

$$ \varphi = \frac{\sigma2 - \sigma3}{\sigma1 - \sigma3} \quad (1) $$

The shape-ratio informs on the stress regime in the region associated with the orientation of the principal stresses (e.g. Gephart and Forsyth, 1984). Equivalently, we will use the R-ratio, defined as $R = 1 - \varphi$:

$$ R = \frac{\sigma1 - \sigma2}{\sigma1 - \sigma3} \quad (2) $$

A R-ratio of 0 indicate a purely extensive regime, while values of 0.5 and 1 indicate strike-slip and compressive regime, respectively. We also calculate the uncertainties associated with the orientation of each stress component and shape-ratio, following the approach of Martínez-Garzón et al. (2016). For the 2014 relocated aftershocks catalogue, we used the mean uncertainty obtained when calculating the focal mechanisms to perturb the strike (+/- 25°), dip (+/- 12°), and rake (+/- 36°). These uncertainties are obtained by calculating 100 times a solution for each focal mechanism. As we
340 did not have information about the uncertainty associated with the published focal mechanism, we 341 chose to perturb the parameters with a maximum of +/- 30°, which is consistent with the 342 uncertainties we obtained for the 2014 relocated aftershocks. Then, randomly perturbing 100 times 343 the strike, dip, and rake of each mechanism assuming a uniform distribution, we create 100 new 344 catalogues, on which we apply 100 bootstraps operations. The inversions then lead to 10,000 values 345 of shape-ratio, trend and plunge for each stress. From then, we calculate the mean and the standard 346 deviation (Tab. S2, Fig. S5, Fig. S6).

347 We then aim at exploring how the stress-state is variable in space and time, calculating it for the 348 different focal mechanism catalogues defined in previous section: full Ubaye Region, Ubaye Valley, 349 Serenne area, Background seismicity, 2003-2004, 2012-2015, and 2014 relocated aftershocks. The 350 results of these inversions are presented in the Results section.

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# 4.2 Fluid-pressure

353 As mentioned in the introduction, fluid-pressure was highlighted as a driving process of the Ubaye 354 Region seismicity. Two previous studies (Daniel et al. 2011; Leclère et al. 2013) calculated the fluid- 355 overpressure needed to explain the 2003-2004 swarm and both assumed a stress-state with two 356 horizontal components (σ1 and σ3) and one vertical component (σ2). However, in our case, the 357 stress-state inversed for the 2012-2015 sequence (Fojtíková and Vavryčuk, 2018) has only one 358 horizontal component (σ3) and two components (σ1 and σ2) neither vertical nor horizontal. 359 Therefore, we generalize the approach used by Collettini and Trippetta (2007), Terakawa et al. 360 (2010), Leclère et al. (2013), and Jansen et al. (2019), to calculate the fluid-overpressure.

361 As neither σ1 nor σ2 are vertical, we first express the vertical stress (σv) as a function of the principal 362 stress components using Cauchy's equation. Assuming a lithostatic stress, we get:

$$
\sigma v = \cos^2(\theta_1)\sigma_1 + \cos^2(\theta_2)\sigma_2 = \rho g Z \quad (3)
$$
With $\theta_1$ and $\theta_2$ corresponding to the angle between the vertical stress and the plunge of $\sigma_1$ and $\sigma_2$. $\rho$ is the rock density taken at 2700 kg/m³; $g$ is the gravitational constant (rounded at 10 m/s²) and $Z$ is the depth in meters.

Using the R-ratio (eq. 2) determined in the stress inversion, we can then reorganize eq. 3 to express $\sigma_1$ and $\sigma_2$ as function of $\sigma_3$:

$$ \sigma_1 = \frac{\rho g Z - R \cos^2(\theta_2) \sigma_3}{(1-R)(\cos^2(\theta_2) + \cos^2(\theta_1))} \quad (4) $$

$$ \sigma_2 = \frac{\rho g Z - R(\rho g Z + \cos^2(\theta_1)\sigma_3)}{\cos^2(\theta_1) + \cos^2(\theta_2) - R\cos^2(\theta_2)} \quad (5) $$

We then assume that $\sigma_3$ depends linearly on the depth ($Z$), such that:

$$ \sigma_3 = AZ \quad (6) $$

Therefore, the only unknown left to determine $\sigma_1$ and $\sigma_2$ in eq. 4 and 5 is the coefficient $A$. To determine it, we first compute the shear and normal stress that apply on the faults. As we do not know yet which of the nodal planes inferred from the mechanism the fault plane is, both nodal planes are considered as possibly activated. The shear stress $\tau$ and normal stress $\sigma_n$ are related to the principal stress components using Cauchy's equations:

$$ \tau^2 = (\sigma_1 - \sigma_2)^2 l^2 m^2 + (\sigma_2 - \sigma_3)^2 m^2 n^2 + (\sigma_3 - \sigma_1)^2 n^2 l^2 \quad (7) $$

$$ \sigma n = l^2 \sigma_1 + m^2 \sigma_2 + n^2 \sigma_3 \quad (8) $$

With $l^2$, $m^2$, and $n^2$, the cosine direction of $\sigma_1$, $\sigma_2$, and $\sigma_3$ respectively.

We can then define the slip-tendency $\alpha$ for all nodal planes following Mattila (2015) and Jansen et al. (2019). This parameter describes how close the fault stress-state is from failure. A fault is at a critical stress state for $\alpha$ equal to 1, while it gets further from failure when $\alpha$ decreases. If we first assume that all faults are at a hydrostatic pressure, $\alpha$ is defined as:
$$ \alpha = \frac{\tau}{\mu(\sigma n - \rho_f g Z)} \quad (9) $$

With $\rho_f=1000$ kg/m³ the water density and $\mu$ the friction coefficient. We here assume a standard value of $\mu=0.6$ (Byerlee, 1970).

When combining eq. 4, eq. 5 and eq. 6 with eq. 7 and eq. 8, the slip-tendency (eq. 9) only depends on A. We then perform a grid search with A varying from 0 to $\rho_g$, with a step of 1 N/m³. For each value of A, we calculate the slip-tendency for all fault planes. We then select the best A-value, 1) given that the slip tendency cannot be higher than 1, and 2) assuming that at least one event occurs with a fluid pressure equal to the hydrostatic pressure, meaning that the most favourably oriented fault plane is reactivated with hydrostatic fluid pressure (Terakawa et al. 2010; Terakawa et al. 2012). This assumption tends to minimize the fluid pressure required to reach the reactivation threshold. Therefore, we are looking for an A-value for which the maximal slip tendency is equal to 1. Once A is determined, $\sigma_1$, $\sigma_2$, and $\sigma_3$ (eq. 4, 5, and 6), as well as the stress state $\tau$ and $\sigma_n$ acting on the faults (eq. 7 and eq. 8) can be computed. Assuming cohesionless fault, we can then compute the fluid-pressure that is required to induce failure using the Amontons' law:

$$ PF = \sigma n - \frac{\tau}{\mu} \quad (10) $$

For each event, a fluid-pressure value is computed for both nodal planes. To discriminate the fault plane from the auxiliary plane, we could use either the misfit angle, which is the difference in the direction between the slip and the maximal shear stress on the fault plane if considering the Wallace and Bott assumption (Wallace, 1951; Bott, 1959), or the slip tendency (eq. 9), hence the lowest fluid pressure, in the Mohr-Coulomb theory. For ~75% of the events, the plane with the highest slip tendency is also the plane with the lower misfit angle or one of two planes with similar misfit angle (see supplementary Fig. S7). For ~25% of the events, there is a discrepancy between the plane with the lower misfit angle and the one with the highest slip tendency. As the (Wallace, 1951) and (Bott, 1959) assumption may show some limits (Kakurina et al., 2019) and because we compute the full
stress tensor, we choose to keep as the fault plane the nodal plane the closest to failure, which is also the plane with the lowest fluid-pressure. The average misfit angles of the chosen fault planes is ~19° for the 2014 aftershocks (see supplementary Fig. S7), with only a few events (7%) with a misfit angle larger than 40°, but lower than 75°. To further validate that our fault plane determination is correct, we compare the chosen fault plane with the one proposed by the fault instability criterion (Vavryčuk et al. 2013). The fault instability criterion gives the nodal plane that is the closest to the rupture without resorting of stress-state computation. It only depends on the friction coefficient μ, the R-ratio, and the direction cosines from the principal stresses (Vavryčuk et al. 2013):

$$I = \frac{\tau_r + \mu (\sigma_r + 1)}{\mu + \sqrt{1 - \mu^2}} \quad (11)$$

where τr and σr are equal to:

$$\tau_r = \sqrt{l^2 + (2R - 1)^2 m^2 + n^2 - (-l^2 + (2R - 1)m^2 + n^2)^2} \quad (12)$$

$$\sigma_r = -l^2 + (2R - 1)m^2 + n^2 \quad (13)$$

We find that 99% of the identified fault planes are similar by both approaches, confirming their determination.

The fluid-pressure depends on the depth as it is assumed to be higher than the hydrostatic pressure: $Ph = \rho_f g Z$ (14). To work with comparable values, we subsequently use the fluid-overpressure, defined as the part of fluid-pressure above the hydrostatic pressure (MPa):

$$\Delta PF = PF - Ph \quad (15)$$

We can then define the overpressure ratio (%) by normalizing the overpressure by the hydrostatic pressure:

$$\frac{\Delta PF}{Ph} = \frac{(PF - Ph)}{Ph} * 100 \quad (16)$$
418 Equivalently, another way to observe overpressure is through the pore-fluid factor (Sibson, 1990; Cox, 2016), which is the ratio of the pore-fluid pressure to the lithostatic pressure. We choose in this paper to use the overpressure ratio to compare with other studies using similar methods (e.g. Jansen et al. 2019), but results are shown in supplementary figures (Fig. S8, Fig. S9, Fig. S10, Fig. S11).

422 At the end, we obtain the fluid-overpressure needed to trigger slip on all fault planes of the focal mechanisms. We also estimate fluid-overpressure uncertainties by randomly perturbing 1000 times the strike, dip, and rake of the focal mechanisms on a range of 30° in a uniform distribution law. We then run the full inversion process on these 1000 perturbed catalogues, to reconstruct the stress-state and overpressures, leading to 1000 perturbed values of fluid-pressure (Fig. S12). Since the distribution is not symmetrical, we quantify the uncertainties for each event by measuring the overpressure range for which 34% (half of one standard deviation if the distribution was Gaussian) of the calculated values are above or below the initial fluid-pressure value. We then compute the median of these uncertainties for the different catalogues. We obtain an overpressure ratio uncertainty of -15% and +21% for Ubaye Valley, of -18% and +24% for Serenne, and of -17% and +15% for 2014 relocated aftershocks.

433

## 5. Results

435

### 5.1 Orientation and amplitude of the stress tensor

436 We determine the stress-state for the two months following the 2014 mainshock as well as for the full Ubaye Region, its two sub-areas (Serenne and Ubaye Valley) and the particular sequences (2003-2004 and 2012-2015 sequences). Fig. 4 summarizes the stress-state orientations calculated here and compares them to the published ones (Leclère et al. 2013; Fojtíková and Vavryčuk, 2018).

439
For the full Ubaye Region, we obtain a σ1 oriented N37° ± 8° and plunging 66° ± 6°, a σ2 oriented N192° ± 2° plunging 22° ± 6°, and a sub-horizontal σ3 oriented N285° ± 1°. The σ3 is the only horizontal stress component and is tilted ~65° from the normal of the Frontal Penninic Thrust. The other two stress components, σ1 and σ2, are less well-resolved (Tab.S2, Fig.S6). Indeed, they can easily vary in a N190° vertical plane, which suggests that both components have similar amplitudes. Furthermore, the Ubaye Region shows strong variations in the topography, with more than 2000 m between valleys and mountain summits. As σ1 and σ2 depend on the depth, topographical effects may impact the plunge, as well as the amplitude, of these stress components.

We then refine this analysis by looking for spatial variations for the two sub-areas: Ubaye Valley and Serenne. In both cases, σ3 orientation is found similar to the Ubaye Region one while there are differences in the plunge for the other two components. Both σ1 and σ2 are within the N190° vertical plane, but σ1 is found subvertical (71°) in the Serenne area, while it shows a 41° plunge in the Ubaye Valley. Accordingly, σ2 is sub-horizontal (14°) and more vertical (49°) in the Serenne and Ubaye Valley, respectively. As mentioned above, this difference might be (partly) due to topographical effects.

The stress-state orientations for the two main sequences of 2003-2004 and 2012-2015 and for the remaining focal mechanisms (background seismicity) in the Ubaye Valley are found very similar to the Ubaye Valley one (Fig. 4, Tab.S2). Some very limited differences (<5°) in the plunge of σ1 and σ2 might be explained by the variation of topography or uncertainties in the computation. Same observations can be noticed when looking specifically at the aftershocks of the 2014 event: the stress-state does not show significant differences with the other ones.

Moreover, the stress-state orientation we determined are similar to the ones computed by Fojtíková and Vavryčuk (2018), the differences being within the uncertainty range. The stress-state computed by Leclère et al. (2013) for the 2003-2004 sequence shows a subvertical σ2, while they use the same catalogue of mechanisms as Fojtíková and Vavryčuk (2018) who found plunging of 49° (Tab.S2).
We also retrieve similar R-ratio ≈ 0.31, for the different catalogues (Tab.S2). Their values are also compatible with the ones found by Leclère et al. (2013) for the 2003-2004 swarm and by Fojtíková and Vavryčuk (2018) for the 2003-2004 and 2012-2015 swarms with a mean R-ratio ≈ 0.29. This value indicates extension and strike-slip regime in this region, which is in agreement with geodetic observations.

<figure><img src="image_6.png" /><figcaption>Figure 4: Polar view of the stress-state orientations (circle for σ1, diamond for σ2, and stars for σ3). Computed stress-states are shown as filled symbols coloured by regions (full Ubaye Region, and its two sub-areas: Serenne and Ubaye Valley, Fig. 2b); or periods (2003-2004 swarm, 2012-2015 sequence, background seismicity, Fig. 2c; and 2014 aftershocks, Fig. 3). The red, pink, and purple unfilled symbols correspond to published stress-states: “F&V” for Fojtíková and Vavryčuk (2018); “L” for Leclère et al. (2013).</figcaption></figure>
478

479 The magnitude of the principal stresses is then determined (Figs 5a, c, and e, for Ubaye Valley,
480 Serenne, and 2014 relocated aftershocks catalogues, respectively). In all cases, they are found to be
481 very similar, as σ3≈15*Z (MPa), with the depth Z in km. Equivalently, the deviatoric stress is found to
482 increase with depth as σ1-σ3≈13*Z (MPa). As expected, σ1 is close to σ2, even if their magnitudes
483 vary slightly among cases, as the stress-state inversion reveals differences, albeit small, in their
484 plunge. Once again, these differences may be explained by topographic effect, as the 2014
485 aftershocks occurred below the Parpaillon massif, which culminates at 3000 m.

486 In summary, the stress-states and the R-ratio are found similar for the different catalogues, and
487 consistent with previously published results. Therefore, anomalies in the stress-state orientation or
488 magnitudes cannot explain the swarm behaviours or the complex aftershock sequences of the 2014
489 earthquake. As the stress-state seems spatially homogeneous constant, we then mainly focus on the
490 Ubaye Valley, Serenne area, and 2014 aftershocks to explore their fluid-overpressure behaviour.
<figure><img src="image_7.png" /><figcaption>Figure 5: Fluid-pressure versus depth (left panels: a, c, e) and histograms of overpressure ratios (right panels: b, d, f), for the Ubaye Valley (a, b), the Serenne area (c, d), and the 2014 aftershocks (e, f). In a), c), and e), the dots show the fluid-pressure required to reactivate the selected fault planes coloured by the slip-tendency at hydrostatic pressure. The depth is below sea level. The light blue, blue, black, and red lines correspond to the hydrostatic pressure, σ3, σ2, and σ1, respectively. In b),</figcaption></figure>
497

d), f), the histograms show the number of events per overpressure ratio (%), while the red lines show the cumulative number of events. The purple and yellow bars correspond to the common negative and positive uncertainty that are found to all the fluid-overpressure values. The pink dashed line shows 80% of the cumulative number of events. The number of earthquakes as function of pore-fluid factor is also available in Supplementary Materials (Fig. S8).

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## 5.2 Fluid-overpressure

504

Figs 5a, c, and e shows the fluid-pressure computed for the nodal planes the closest to rupture. These planes are assumed to be the fault planes of the seismic event. This assumption is supported by the correlation of our results with the values of the fault instability criterion described in section 4.2. All computed fluid-pressures are equal or above the hydrostatic pressure, as we assumed a minimal hydrostatic pressure. Only a few events require an overpressure higher than σ3. In this case, failure might imply some tensile component. Some of the published focal mechanisms used in our catalogues have quality indexes (Mathey et al. 2021a). We find that events with a fluid-pressure value higher than σ3 have the lowest quality index (Figs S11b and d). We then discard events with a fluid-pressure value higher than σ3 from the analysis as their mechanisms might be fraudulent with large errors.

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To get rid of the dependency on depth, we also look at the overpressure ratios. We consider that the overpressure required to trigger the events is significant when it is higher than the computed minimum uncertainties: 15% for Ubaye Valley (Fig. 5b), 18% for Serenne (Fig. 5d), and 17% for the 2014 aftershocks (Fig. 5f). Considering these uncertainties, 68% (Ubaye Valley), 61% (Serenne), and 69% (2014 aftershocks) of the events required overpressure to be triggered, leading to median overpressure ratio of 27%, 30%, and 32% for the Ubaye Valley, Serenne, and 2014 aftershocks, respectively. In absolute values, at least 46% of the events require fluid-overpressure less than 45
521 MPa for the Serenne area. For the 2014 aftershocks catalogue, at least 60% of the events need less
522 than 27 MPa of overpressure. 48% of events require less than 38 MPa of overpressure for the case of
523 the Ubaye Valley area (Fig. S14). The three catalogues lead to similar values of median overpressure:
524 24 MPa, Ubaye Valley catalogue; 30 MPa, Serenne catalogue; 20 MPa, 2014 relocated aftershocks
525 catalogue.

<figure><img src="image_8.png" /><figcaption>Figure 6: Map view of the overpressure ratio (%) for the Ubaye Valley (a), the Serenne area (b), and the 2014 relocated aftershocks (c). In c), the red star shows the location of the 2014 mainshock</figcaption></figure>
(Thouvenot et al. 2016). The dashed red-lines show the cross-sections that are used in d) and e). In c), d), and e), the hexagrams and circles correspond to events belonging to swarm-like clusters and aftershock-like clusters, respectively. The depth is below sea level. The same figure, with pore-fluid factor as function of location is available in Supplementary Materials (Fig. S9).

We then analyse the spatial distribution of the reconstructed fluid-pressure (Fig. 6), but we do not see any clear tendency. Indeed, despite an important scattering in overpressure ratios, no consistent anomalies are observed on map view, neither for the 2014 relocated aftershocks (Fig. 6c) nor for the Ubaye Valley and Serenne areas (Figs 6a and b). The families that composed the 2014 aftershocks were classified by De Barros et al. (2019) into aftershock-like or swarm-like behavior based on their space, time and magnitude distribution (Figs 6c, d, and e). The median value of overpressure ratios for aftershock-like clusters is slightly lower (27%) than the one for swarm-like clusters (31%). This is to be expected, as swarms are commonly driven by fluid-driven processes. However, no clear difference is observed spatially between events belonging to clusters with a different behavior (Figs 6c, d, and e).

There is also no clear dependency between the depth and fluid-overpressure for the Ubaye Valley (Fig. 7a), as the median values are found nearly constant between 4-12km depth. For the 2014 aftershocks, we also observe a constant fluid-overpressure as a function of depth (Fig. 7c). Therefore, for both catalogues, the overpressure does not seem to depend on the hydrostatic pressure. On the contrary, for the Serenne area, the reconstructed overpressure seems to increase with depth, and therefore with the value of the hydrostatic level. This difference in behaviour between Serenne and Ubaye Valley might suggest a difference either in the geological context or in the driving processes.
<figure><img src="image_9.png" /><figcaption>Figure 7: Fluid-overpressure (MPa) versus depth for the Ubaye Valley (a), the Serenne area (b), and the 2014 relocated aftershocks (c). Each dot corresponds to a focal mechanism and depth is directly associated to each solution. Each dot corresponds to an event with a computed focal mechanism. Red stars show the median value of the fluid-overpressure on a 500-m-depth sliding window with a 250 m overlap. Only median values calculated with at least 5 events are represented. Empty circles correspond to events having an overpressure above σ3, and are thus discarded for the median calculation. The depth is below sea level.</figcaption></figure>

Fig. 8 represents the overpressure ratio as a function of the event magnitude (ML) for the different areas and seismic sequences. For the sake of homogeneity, we use the local magnitude (ML) computed by the LDG (Duverger *et al.* 2021) for all events in our focal mechanism catalogues. In both
the Ubaye Valley and the Serenne areas, the median of the reconstructed fluid-pressure values does not depend on the event magnitude. The same is observed for the overpressure (Fig. S15). Particularly, the largest events need similar level of fluid-pressure as the average for the smallest ones. The same observation is made when focusing on particular sequences, such as the seismicity in 2012 (Fig. 8d). The reconstructed overpressure ratio appears more scattered for low magnitude events (ML<2.5) as the number of events decreases with magnitude. This might also indicate a variability in the processes or a larger uncertainty on the focal mechanisms for lower magnitude events. However, the mainshock that started this sequence (ML4.8, 26/02/2012) needs ~30% of overpressure ratio (overpressure of ~35 MPa, Fig. S15d), which is above the median value of the reconstructed overpressure ratio and one of the largest for events of magnitude larger than 2.5 (Fig 8d). This suggests that fluid-pressure is required to trigger both large and small events, whatever their magnitude.

To confirm this statement, we look in detail at fluid-pressure uncertainties for particular earthquakes. Figure 9 represents the distribution of fluid-overpressures obtained by perturbing 1000 times the source fault properties (strike, dip, and rake) of all events (see section 4.2). For each of these perturbed catalogues, we run again the full analysis, including the inversion for the stress-state orientation, amplitude and the fluid-overpressure. We then focus on 9 events with magnitude ML from 2.9 to 5.2. Two of these earthquakes occurred in the Serenne area (Figs 9b and c) in Feb. 1996, and 2000, two were from the background seismicity of the Ubaye Valley (Figs 9a and f) in Jan. 1996 and 2005, and two happened during the 2003-2004 swarm (Figs 9d and e). The last three events are the largest earthquakes of the 2012-2015 sequence, which occurred on 26/02/2012 (Fig. 9g), 07/04/2014 (Fig. 9h), and 06/11/2015 (Fig. 9i). Some of these events unambiguously need fluid-overpressure to reach the reactivation criterion (Figs 9b, c, e) as their distributions show a clear peak around a high fluid-overpressure value. The overpressure distributions of the earthquakes occurring in 2003 and 2012, (Figs 9d, g and h) are wider but still suggest an activation by fluid-pressure. However, this is not as clear for the Jan. 1996, 2009 and 2015 events (Figs 9a, f and i) for which their
distribution is squeezed toward low values of overpressure. Overall, by considering uncertainties, the data presented in Fig. 9 confirm the need for overpressures triggering earthquake, even if the reconstructed values of overpressures might be variable. Therefore, this confirms that even the largest earthquakes are likely to have been caused by fluid pressure, whether they occurred as part of a swarm or as an isolated earthquake. To refine this observation, we look at how the overpressure ratio evolves with time within a swarm and during an aftershock sequence (Fig. 10). During the 2003-2004 swarm (Fig. 10a), the overpressure ratio varies in time at high level without following a clear pattern. On the contrary, during a mainshock-aftershocks sequence, like in 2012 (Fig. 10b), high values of overpressure ratio (up to 68%) are observed at the beginning of the sequence, right after the mainshock. It then slowly decays in the next 6 months, to stabilize below 30% (Fig. 10b, black line). Therefore, it suggests that the difference between a swarm and a mainshock-aftershock sequence may be more related to the evolution of the fluid-overpressure in time rather than to its absolute values. Both required overpressure in their driving process, but the overpressure ratio decays with time in the mainshock-aftershocks sequence and not in a swarm episode.
<figure><img src="image_10.png" /><figcaption>Figure 8: Overpressure ratio as function of local magnitude for the Ubaye Valley (a), the Serenne area (b), the 2003-2004 swarm (c), and after the 2012 mainshock (d). The local magnitude comes from the LDG catalogue (http://www-dase.cea.fr). The empty circles show events with overpressure ratio higher than σ3, which are therefore discarded from the analysis. The black line shows the floating median calculated for every 0.25 bin of magnitude with at least 5 events. The same figure, with pore-fluid factor as function of magnitude is available in Supplementary Materials (Fig. S10).</figcaption></figure>
<figure><img src="image_11.png" /></figure>

Figure 9: Distribution of overpressure obtained by the inversion of 1000 perturbed catalogues of focal mechanisms. For each perturbed catalogue, the stress-state is first inverted before being used in the pressure calculation (see section 4). The number of realizations is the number of perturbed catalogues leading to fluid-overpressure value in a 1 MPa bin. Results are shown for 9 selected earthquakes occurring in the background seismicity (a, f), along the Serenne fault (b, c), during the 2003-2004 swarm (d, e) during the 2012-2015 sequence (g to i). The local magnitude of the events is taken from the LDG catalogue (http://www-dase.cea.fr). The dashed-red line corresponds to the initial value of fluid-overpressure calculated using the unperturbed catalogue.
<figure><img src="image_12.png" /></figure>

Figure 10: Overpressure ratio as function of time for the 2003-2004 swarm (a) and after the 2012 mainshock (b). The colour corresponds to the local magnitude from the LDG (Duverger et al. 2021). The grey bars (right axis) show the temporal distribution of the earthquakes, from the LDG catalogue (http://www-dase.cea.fr). Only the events with an overpressure ratio lower than σ3 and in the LDG catalogue are represented. In b), the black line fits the 90th percentile maximum overpressure ratio value every month following an exponential law. The same figure, with pore-fluid factor as function of time is available in Supplementary Materials (Fig. S11).

# 6. Discussion

From the focal mechanisms, we reconstructed the stress-state and the fluid-overpressure that trigger earthquakes for different areas and time periods in the Ubaye Region. We showed that the majority of the events need overpressure at median value of 20-30 MPa, or equivalently at 27-32% of the hydrostatic pressure. We also showed that there are no significant spatial variations of either the stress-state nor the fluid-overpressure. Even the largest events are triggered by fluid-pressure, but the time evolution of fluid-overpressure is different for a mainshock-aftershocks sequence or for a swarm sequence.
641

In this discussion, we first analyse the different sources of uncertainties that may impact our results.

642

Then, we discuss the stress state and overpressure results, in order to propose an interpretative
mechanism for fluid-driven aftershock seismicity and swarms.

643

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645

**6.1 Source of uncertainties**

646

The main uncertainties we have to deal with come from the accuracy of the focal mechanisms, as they are used to compute both the stress state and the overpressures. However, we do not have precise information on the uncertainties of the focal mechanisms we used, as they come from various origins. We therefore have to assume a range of uncertainties. Then, the statistical approach based on the focal mechanism’s perturbation (section 3.2) allows the propagation of those uncertainties through the full workflow to estimate the stress-state and fluid-pressure uncertainties.

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The error bars in Fig. 5 or the distribution of overpressures for the 1000 perturbed catalogues (Fig. 9) show that the reconstructed overpressures may vary, but that the requirement of fluid pressure is reliable.

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We show that the reconstructed fluid-pressures, on average, do not show horizontal spatial variations (Fig. 6). On average, they are also found constant with depth (Fig. 7) in the Ubaye Valley, while they seem to increase with depth in the Serenne area. These observations are based on our confidence in the location of events. However, we do not have information on location uncertainties the focal mechanisms for the regional catalogues. On the contrary, the 2014 aftershocks were relocated through a double-difference algorithm, with limited relative location uncertainties (De Barros *et al.* 2019). Using this specific catalogue, we do not observe any horizontal spatial pattern (Fig. 6c) but a constant trend with depth (Fig. 7c), similarly to the Ubaye Valley patterns. By extrapolation, we may conclude that, on average, the location uncertainties do not likely bias the spatial results.

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664
In our analysis, we focused only on the variation of fluid-overpressure (ΔPF). However, to reach failure, we could also have an increase of shear stress (overstress, Δτ). Particularly, the aftershocks that are close from the 2014 mainshock can be reactivated by Coulomb stress transfer from the mainshock (Stein, 1999). In this case, the overstress should be smaller than the stress drop of the mainshock, which is assumed to be close to 10 MPa (Kanamori and Anderson, 1975). From eq. 10, the required overstress Δτ can be expressed as a function of the computed ΔPF:

$$ \Delta\tau = \mu\Delta PF \quad (17) $$

For most of the events (~68%), the overstress that is needed is above 10 MPa. Therefore, even if part of the required perturbation might be due to shear-stress increase, the reconstructed values suggest that fluid-overpressure is anyway required for a large majority of events. As we could not discriminate between shear-stress increase and fluid-overpressure, and for sake of simplicity, we therefore focus our interpretations on fluid-driven processes.

## 6.2 Stress-state

The reconstructed stress-state shows a horizontal σ3, striking ~N280°, and σ1 and σ2 in a vertical plane oriented ~N190°. Their plunge in this plane varies, as σ2 is found between sub-horizontal and 45° dipping. Their reconstructed stress magnitude is however very close (Fig. 5), meaning that the plunge variations can be easily explained, by, for example, topographical effect. In NE Japan, beneath low-altitude areas, the reverse regime (σ3 vertical) is dominant while high altitude areas show a strike-slip regime (σ2 vertical) (Yoshida et al. 2015). Similarly, the Ubaye Region shows large altitude differences (~2000m) between summits and valleys. Therefore, additional weight from high summit mountains may increase the vertical stress, leading to a σ1 plunge closer to the vertical beneath summits, while it will get more horizontal beneath valleys.
687 The inversion of focal mechanisms for the different sub-areas lead to very similar stress-state, in both
688 its orientation and its amplitude. The stress-state orientation is consistent with results from other
689 studies (Fojtíková and Vavryčuk, 2018). Particularly, similar stress-states are found for the Serenne
690 area, characterized by well-identified faults and diffuse seismicity, and for the Ubaye Valley where
691 recent seismicity occurred as swarms or mixed sequence along hidden faults. Therefore, large spatial
692 heterogeneity in the stress-state cannot be used to explain the variability in the seismic behaviours,
693 even if we cannot discard very localized variations (Heap et al. 2010) that are below the resolution
694 we have with focal mechanisms inversion.

695 We do not observe differences in the stress-state either when using the full time period or focusing
696 on perturbed periods of dense seismic activity, such as during the 2014 aftershocks sequence. We
697 could have expected a perturbation of the stress-state because of aseismic afterslip that drives the
698 aftershocks. For example, Duboeuf et al. (2021) observed perturbed stress-state associated with fault
699 reactivation by fluid-injection in decametric experiments. As more than 90% of the deformation was
700 observed to be aseismic, they attributed the stress perturbation to the aseismic slip. Same conclusion
701 was reached by Schoenball et al. (2014) during a geothermal exploitation. Unlike aseismic slip, fluid-
702 pressure may not influence the determination of stress-state orientation. Indeed, as fluid-pressure
703 could trigger slips on faults that are badly oriented in whatever direction toward the stress-state, the
704 focal mechanisms inversion leads to an average stress-state, similar to an unperturbed one.
705 Therefore, the constant stress-state we reconstructed suggests that large aseismic slip is not a likely
706 mechanism to explain the high rate of seismicity. Even if the spatial resolution brought by the focal
707 mechanisms cannot allow us to conclude on the presence of aseismic slip at small-scale, fluid-
708 pressure seems to be the dominant process.

709

710

## 6.3 Fluid-pressure values
Several studies (Jenatton et al. 2007; Daniel et al. 2011; Leclère et al. 2013; Thouvenot et al. 2016; De Barros et al. 2019) highlighted the role of fluid-pressure as the main driving process required to explain the high rate of seismicity for the 2003-2004 swarm and the 2012-2015 aftershocks. We here find that significant overpressure (above uncertainties) is required for most of the events (~77%), with a median overpressure of 22 %. This observation is independent of the areas, as both Serenne area, and Ubaye Valley area, show similar pressure levels, despite their different seismic behaviour. This need for fluid-pressure is also seen for all types of seismic sequences, including mainshocks, their following aftershocks, or swarm sequences (Figs 6d and e). On average, we find that the overpressure ratio should be up to 40% and, equivalently, that the fluid-overpressure is up to 70 MPa, even if 80% of the events need less than 40 MPa. Other fluid-overpressure analysis was carried out in the Ubaye Valley for the 2003-2004 swarm by Leclère et al. (2013). Using a Mohr-Coulomb analysis, they obtained a required overpressure of up to 55 MPa, consistently with the ones we found for the 2003-2004 period and for the other periods and areas.

Determination of fluid-overpressure was also done for other swarms in different contexts. The Mogul swarm (2008, e.g. Ruhl et al. 2016) occurred in the Nevada state in a hydrothermal context. Using a similar methodology as the one used here, Jansen et al. (2019) computed the fluid-overpressure required to trigger the Mogul seismicity and confirmed their results with numerical modelling. They obtained a median overpressure ratio of 42% that is more than for the full Ubaye Region (22%). This difference might be linked to the presence of hot springs in the Mogul swarm area, which revealed an elevated fluid-pressure level in the medium. They also found fluid-overpressure up to 30 MPa, which is consistent with our values. For some events, they also found fluid-pressure higher than the minimum magnitude stress component. They proposed that other processes (static stress transfer, transient poroelastic stress changes, aseismic slip) should explain the high fluid-pressure values obtained.
Fluid-driven process was also proposed for the Vogtland/Western Bohemia area that was struck by several swarms (e.g. Heinze et al., 2017). Kurz et al. (2004), using finite element modelling, calculated the fluid-overpressure required to trigger the different earthquake swarms in the 1990-1997 period. They estimated a need of fluid-overpressure up to 50 MPa, with some events needing overpressure as high as 150 MPa, which they deemed unrealistic. Hainzl et al. (2012) also computed fluid-overpressure for the beginning of the 2008 Western Bohemia swarm. Using stress changes and hydrofracture model, they estimated a need of 20-30 MPa of fluid-overpressure. Terakawa et al. (2010) used a similar method as Leclère et al. (2013) to estimate the fluid-overpressure required for the Mw6.3 L'Aquila earthquake and its foreshocks-aftershocks. They found a range of fluid-overpressure between 20 to 50 MPa. In this case, even if a strong earthquake was likely induced by a fluid-driven process, the range of fluid-overpressure is similar to the one inferred on natural swarms where the magnitude rarely exceeds ML4 (Ma and Eaton, 2009). Therefore, it seems that, for natural swarms induced by fluids, fluid-overpressure with values up to ~50 MPa is observed.

During injection-induced swarms, the behaviour of the fluid-pressure and therefore the fluid-overpressure is monitored. Baisch et al. (2010) observed the evolution of the fluid-pressure in a case of fluid injection experiments in Soultz-sous-Forêts (France). They observed a fluid-overpressure up to 18 MPa at the reservoir depth (~4km). For another experimental zone, Basel Switzerland, Terakawa et al. (2012) retrieved a required fluid-overpressure up to 30 MPa. These values of fluid-overpressure in injection-induced swarms are also similar to the ones we found here and for other natural swarms (Mogul, Vogtland/Western Bohemia).

## 6.4 Origin of pressurised fluids

One of the main questions that remains when involving fluids in the driving processes is where they come from and how they get pressurized. A first possibility is the meteoric origin. Some swarm
sequences (e.g. Mt Hochstaufen, Hainzl et al. 2006; Mogul, Jansen et al. 2019) were explained by the infiltration of rainwater in the medium that increased the fluid-pressure and triggered the seismicity.

For the 2003-2004 swarm, Leclère et al. (2012, 2013) also proposed that meteoric fluids infiltrated in the Argentera massif and then flowed to the seismic area of the 2003-2004 swarm. The difference in elevation between the Argentera summits (>2500 m) and the Ubaye Valley (~1000 m) allow a fluid-pressure to develop in the swarm area, assuming that faults are highly permeable with the Embrunais-Ubaye nappes acting as cap-rock. This is consistent with the chemical analysis of a geothermal spring located 20 km south-east of the Ubaye Valley (Bagni di Vinadio, Baietto et al. 2009; Perello et al. 2001). Water is shown to have a meteoric origin, before flowing up to ~5km depth, close to the seismicity depth. Such a model was also used to explain the seismicity of the Mogul swarm, which also occurred below a valley surrounded by high summits (Ruhl et al. 2016; Jansen et al. 2019). Following this idea, we could expect to see a spatial tendency in the overpressure distribution, with a decrease of overpressure with the distance to the Argentera and Pelvoux summits. This is not observed, as the overpressure does not show any clear spatial tendency.

Moreover, as such a model requires a cap-rock to allow pressure to develop, the seismicity away from the Embrunais-Ubaye nappes, as the Serenne earthquakes, should not require fluid-pressure.

Finally, if the Embrunais-Ubaye nappes acted as a cap-rock, the seismicity would dominantly develop just below it, at ~2 km depth below sea level. This is not observed, as the seismicity mainly equally develops from ~4 to 12 km below sea level.

Another possibility is the presence of a deep reservoir. This was proposed to explain the seismicity of several swarms (Duverger et al. 2015, 2018; Hainzl et al. 2016; Jansen et al. 2019; De Barros et al. 2020; Ross et al. 2020). The reservoir is at first impermeable, then due to some perturbations (stress-change, tectonic loading ...), it becomes permeable by place, which releases the fluids that then flow in the medium. The fluids could migrate either upward (Duverger et al. 2018; Jansen et al. 2019) or horizontally (Farrell et al. 2010). This is for example the main interpretation for the complex seismic behaviours in the Corinth rift. The low-permeability Phyllade nappe at 6-8 km depth leads to
pressurized fluid beneath it. Its failure then allows fluid-pressure to diffuse, triggering swarm seismicity (Duverger et al. 2018). In the Ubaye Valley, the reservoir needs to be of the scale of the studied region, as fluid-pressure is required for the majority of the events. This might not be likely, as, in this case, overpressures are expected to be at similar values and not scattered as observed. Moreover, a hydrological barrier at depth is required to maintain the reservoir at high pressure, but there is no geological evidence for such a layer at a regional scale. However, a smaller reservoir beneath the starting location of the 2003-2004 swarms could explain it, as seismic migrations might then be related to fluid flow (Shapiro et al. 1997).

A third possibility to explain the fluid-pressure is that the fluids are locally pressurized within the faults, through a model of fault-valve behaviour (Sibson, 1990) or creep-compaction (Leclère et al. 2015). If the fault is sealed by a hydraulic barrier with low-permeability, the fluid-pressure tends to increase from a hydrostatic pressure to a lithostatic one. This increase of pressure then promotes the fault slip. Doing so, the hydraulic barrier in the fault may break, causing the over-pressurized fluids to migrate upward. This model was used to explain different swarms (e.g. Cahuilla swarm, Ross et al. 2020; Southern California, Ross and Cochran, 2021; Haenam swarm, Son et al. 2021; Mammoth Mountain, Shelly et al. 2015). A remaining question is the nature of the hydraulic barrier within the fault. There is so far no evidence of a layer in the basement at depth that could produce a low-permeability cap-rock. Leclère et al. (2015) suggested that creep compaction of the fault can lead to impermeable barriers in the fault zone. Alternatively, the presence of clay within the faults by direct chemical and physical alteration may produce hydraulic barriers. Finally, the presence of Embrunais-Ubaye nappes, which lie on clay-rich materials (Lias, Trias, and Callovo-Oxfordien) may prevent a fast overpressure dissipation, which may explain why seismicity organizes as swarms in the Ubaye Valley and not in the crystalline massifs around (Argentera and Pelvoux). We therefore detail how such a model may be applied in the Ubaye Valley in the next section.

809
810

811

## 6.5 Mechanisms for fluid-driven aftershock seismicity and swarms

812

813

814

815

816

817

818

819

820

821

822

823

824

825

826

827

828

829

830

831

832

The fault valve behaviour is the model we use to explain the different types of seismicity, and particularly the high rate of aftershocks observed after the 2014 events. As the required fluid-pressure decreases with time, the mainshock is likely to initiate the fluid-pressure diffusion. Initially, the fluids are trapped by hydraulic barriers, made of clay lenses from gouge materials or mineral recrystallization, within the 2014 mainshock fault as it may tend toward lithostatic pressure (Fig. 11, Initial State). The fluid-pressure then increases as it equilibrates with the confining stress through time, and may tend toward the lithostatic pressure. This leads the faults reaching the reactivation criterion, generating the 2014 mainshock, which also breaks the hydraulic asperities. The pressure then readjusts along the rupture plane, and trigger events on smaller faults close to or above the main fault plane (Fig. 11, Perturbed State). This is consistent with a fluid-overpressure constant with depth, as observed in Fig. 7c. The fluid-pressure, together with Coulomb stress transfer, therefore participates in the triggering of the aftershocks close to the main fault plane, as fluid-pressure seems also required for some of the deep events around the mainshock (Fig. 7c). As this mechanism mainly involved fluid-pressure readjustment, with limited fluid flow, it does not imply a seismic migration with depth, which indeed was not observed in the 2014 aftershocks. Later on, in the shallow clusters, fluids flow along the reactivated planes leading to seismic migrations consistent with the observations made by De Barros et al. (2019) (Fig. 11, Perturbed state). By this mechanism, the pressure decreases and tends to equilibrate with hydrostatic pressure, leading to the decay with time of the reconstructed pressure. Globally, the very high rate of aftershocks observed during this sequence may be explained by the release of overpressure by the mainshock on top of the coseismic stress transfer.

833

834

During a swarm sequence, like in 2003-2004, the fluid-pressure does not decrease with time (Fig. 8a) and triggered seismicity for nearly two years. Moreover, seismic migration consistent with fluid
diffusion is observed for swarms. Therefore, a source of fluid that continuously feeds the swarm, such as a reservoir, is likely beneath the swarm. By analogy with the mainshock-aftershock sequence, the reservoir could be a set of fractures and faults in which fluids are initially trapped and pressurized, even if our study does not bring any information on its nature. However, the initial failures that then release the fluids are probably aseismic here, likely because they are at large depth, below the seismogenic depth. Creep compaction associated with an aseismic slip at depth (Leclère et al. 2015) might also be a mechanism to pressurize the fluids. In addition to pressure perturbation, fluid is flowing, leading to seismic migration. In this case, the swarm sequence corresponds to the last step of the model we proposed for the mainshock-aftershocks sequence, similarly to the swarm behaviour sequences observed in the shallow clusters in 2014 (Fig. 11, Perturbed State, top). In the 2003-2004 case, a large volume of water is supposed to flow in order to maintain the swarm active for two years with a seismic migration. The volume that triggers this swarm is estimated to be around 1e7 m³ (Danré et al., 2022b), similarly to e.g. the long duration Cahuilla swarm, which was also interpreted to be driven by natural fluid-injection (Ross and Cochran, 2021). Assuming a reservoir made of a fracture network with a similar size as the seismic swarm, its total volume will be about 32 km³ (8 km x1 km x4 km; Daniel et al., 2011). Therefore, the fluid volume (1e7 m³) required to induce this swarms may be reached by an equivalent porosity decrease of only 0.03%. Such a value seems reasonable considering the porosity of a fractured fault zone that may act as reservoir (Géraud, Diraison and Orellana, 2006; Jeppson, Bradbury and Evans, 2010; Jeanne et al., 2012). However, this process also needs to be combined with the shear-stress transfer. Indeed, as mentioned by Daniel et al. (2011), the earthquake-earthquake interaction triggered 59% of the seismicity for the 2003-2004 swarm, and the 41% remaining then might be associated with the fluid-pressure processes. This model may also explain the 1977-1978 swarm (Fréchet and Pavoni, 1979) and the 1989 swarm (Guyoton et al. 1990) that occurred in the Serenne area, indicating similar processes in this area.
860 In the Serenne area, fluid-overpressure is found increasing with depth, as does the hydrostatic
861 pressure. Therefore, it is likely that the pressurization occurred at different depths within the faults.

862 In our model, the seismic behaviours observed there can be similar to mainshock-aftershocks
863 sequence, where the mainshock is triggered by fluids and breaks barriers to allow pressure to diffuse
864 upward. The pressure might then trigger the aftershock seismicity, similarly to the 2014 mainshock
865 process. Moreover, in the Serenne area, the seismicity is more diffuse than in the Ubaye Valley. It
866 might mean that 1) the fluids might be locally pressurized within the faults at different depths, 2)
867 there is no fluid-pressure feeding system at depth that can trigger large swarms, and 3) fault areas
868 that can be pressurized are of smaller size or disconnected. Therefore, the main differences between
869 these two zones may be related to its fault architecture, and particularly to its hydraulic barriers.

870 Finally, the differences in behaviour between Ubaye Valley and Serenne area might be also impacted
871 by the differences in the lithology composition and by the presence of the Ubaye nappes. In
872 particular, the latter, by acting as a shallow cap-rock, may lead to a slow decay pressure
873 perturbation, hence to much longer seismic sequences in the Ubaye Valley.

874
<figure><img src="image_13.png" /><figcaption>Figure 11: Conceptual model to explain the different behaviours of the seismicity in the Ubaye Valley.</figcaption></figure>

Left, initial state without medium perturbations other than tectonic loading. Right, perturbed state showing mainshock-aftershocks behaviour (bottom) and swarm behaviour (top).

## 7. Conclusion

Based on the reconstructed focal mechanisms, we determined the stress-state and estimated the overpressures required to trigger earthquakes in the Ubaye Region. We refined this analysis by focusing on specific areas and periods of seismic activity. For the Ubaye Valley, the average stress component orientations are N24°E, 41° for σ1, N201°E, 49° for σ2, and N293°E, sub-horizontal for σ3.
Similar values were obtained for the Serenne area, with some differences in the plunge of σ1 and σ2,
which may be due to topographic effects.

We found that most of the events (77%) are triggered by fluids, with a median overpressure of 22%.
Overall, the majority of the events (65%), for all sub-areas and periods, required between ~17% and
~40% (~15 MPa and ~40 MPa) of overpressure, which is similar to other natural or induced swarms.
Even the events with the largest magnitude seem to be triggered by fluid-overpressure. Therefore,
mainshock-aftershocks sequences and swarms are both likely triggered by fluids. The difference in
behaviours do not seem to be linked to the overpressure values, but to its time evolution. While
overpressure decays with time in an aftershock sequence, it stays at a high level for swarms. Hence,
different origin of fluid-pressure may be at play. For the mainshock-aftershock sequences, we
proposed that the main faults get pressurized, which trigger the first, largest earthquake. Seismicity
is generated on shallower structures through fault-valve behaviour and stress balance. This model is
in accordance with the observations made by De Barros et al. (2019). On the contrary, swarms need
a constant feeding of pressure from a deeper reservoir, likely below the seismogenic depth.
Therefore, we here highlighted the role of fluid in both types of earthquake sequences.

902

903

904 Acknowledgements

Some of the figures were made using GMT (Wessel *et al.* 2019), and the colour scales used in the different figures come from the Scientific Colour Map of Crameri (2018). The focal mechanism catalogues are taken from the supplementary materials of Thouvenot *et al.* (2016), Mathey *et al.* (2021a), and Mazzotti *et al.* (2021).

M. Baques would like to thank IRSN and CEA for her funding grants. We would also like to thank
Henri Leclère, which provides the focal mechanisms catalogue for the 2003-2004 swarm. We also
911 thank the CEA, which provides us the LDG catalogue for the 1965-2020 period in the studied area.

912 We thank the editors, two anonymous reviewers and F. Thouvenot for their thoughtful comments

913 that have helped to improve the manuscript.

914 **Data Availability**

915 Seismic data used for the focal mechanisms of the 2014 aftershocks are available online through the

916 RESIF portal (http://seismology.resif.fr/), with permanent stations on

917 https://doi.org/10.15778/RESIF.FR and temporary ones on https://doi.org/10.15778/RESIF.XG2014

918 and through the INGV portal (http://terremoti.ingv.it/), with permanent stations on

919 https://doi.org/10.13127/SD/X0FXNH7QFY. The resulting focal mechanisms catalogues, the stress-

920 states, the fluid-pressure values, and the codes developed during this paper are available in request.

921 The LDG catalogue (http://www-dase.cea.fr) used in this study is provided in the supplementary

922 materials.

923 **Supporting information**

924 Supplementary figures and data:

925 Supplementary_materials.docx

926 LDG catalogue.txt

927

928

929

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