Maria Sachpazi

Vasilis Kapetanidis

Marinos Charalampakis

Mireille Laigle

Edi Kissling

Anna Fokaefs

Elena Daskalaki

Ernst Flueh

Alfred Hirn

Methoni Mw 6.8 rupture and aftershocks distribution from a dense array of OBS and land seismometers, offshore SW Hellenic subduction

Keywords: Methoni earthquake, SW Hellenic subduction, Aftershocks, Afterslip, Upper plate deformation

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Introduction

The southwestern offshore megathrust domain of the Hellenic subduction has hosted the largest earthquake of the Mediterranean area, a magnitude 8-8.5 tsunamigenic event in 365 AD [START_REF] Papazachos | The Earthquakes of Greece[END_REF][START_REF] Stiros | Seismicity of Western Crete and the destruction of the town of Kisamos at AD 365: archaeological evidence[END_REF]). Yet, only five interplate events with moderate magnitudes (~6.5-7.0) occurred during the instrumental period (1965-present) (Fig. 1a), and only three such events occurred during the earlier period in the 19 th to early 20 th century [START_REF] Papazachos | The Earthquakes of Greece[END_REF].

On 14 February 2008, a Mw = 6.8 subduction thrust event occurred, 50 km offshore of the southwest coast of Peloponnesus, known as the Methoni earthquake (Fig. 1a). Despite its moderate magnitude, it is the largest event, for half a century, in the 400 km long part of the Hellenic subduction, from the Cephalonia Transform Fault CTF) to Crete. The Methoni event ruptured a ~ 30 km x 30 km thrust area along the plate interface between the southwestward overriding Aegean plate and the Ionian oceanic crust [START_REF] Roumelioti | The 14 February 2008 earthquake (M6.7) sequence offshore south Peloponnese (Greece): Source models of the three strongest events[END_REF]. Geodetic measurements (continuous GPS time series) have shown that post-seismic slip continued for ~3 years and the total displacement is of comparable magnitude as co-seismic slip (Howell et al., 2017a). The post-seismic displacements are attributed to afterslip on the subduction interface, distributed over a 100kmX120 km wide region including the co-seismic rupture with its greatest values observed where elevated levels of seismicity were detected following the earthquake (Howell et al., 2017a) (Fig. 1b).

Afterslip occurs as a widespread post seismic process in subduction zones, complementary to the large coseismic slip zone (e.g. [START_REF] Cattania | Aftershock triggering by postseismic stresses: A study based on Coulomb rate-and-state models[END_REF]Yagi et al., 2003;[START_REF] Ozawa | Preceding, coseismic, and postseismic slips of the 2011 Tohoku earthquake, Japan[END_REF] and it is considered as playing an important role in promoting aftershock earthquakes on the subduction interface [START_REF] Hsu | Frictional afterslip following the 2005 Nias-Simeulue earthquake, Sumatra[END_REF] and the upper plate as well (Hayes et al., 2014a).

In this paper, we present observations collected from 9 to 21 months after the Methoni earthquake by a dense array of Ocean Bottom Seismometers (OBS) and seismic stations in the close onshore. We investigate post-Methoni seismicity with respect to the shape and extension of the rupture area as well as the widely distributed post seismic deformation. We combine seismic activity, focal mechanisms and structural data to characterize the pre-, co-and post-seismic behavior of the megathrust boundary over the Methoni rupture area, as well as the regional seismicity. Methoni earthquake location (indicated by its CMT Harvard focal mechanism beachball) is from [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. Focal mechanisms for subduction thrust events of Mw 6.0 or greater during the instrumental period are from the CMT Harvard catalogue, except for the Crete 2013 event (Howell et al., 2017a). Earthquake locations of these events are from the National Observatory of Athens (NOA) catalogue. Here and in subsequent figures: Double-lines: Slab along-dip faults from Sachpazi et al. (2016a), labelled F5 to F9. Dashed double lines: offshore extensions of mapped onshore slab faults from [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. Colored lines: isobaths of slab Moho with depth denoted in kilometres (Sachpazi et al., 2016a,b). NMT and SMT: the Northern and Southern Matapan Trough. Black barbed line: outer limit of backstop, thrust contact of the accretionary wedge of the Mediterranean Ridge over it. Black dashed NW-SE trending line: Inner limit of the backstop at the transition between the Matapan Trough to the SW and the Hellenic continental margin to the NE from Le [START_REF] Le Pichon | A new approach to the opening of the Eastern Mediterranean Sea and the origin of the Hellenic Subduction Zone. Part 2 The Hellenic Subduction[END_REF]. (b) Distribution of post-seismic slip over the offshore forearc as afterslip according to geodetic data modeling (Howell et al., 2017a) with the higher >0.2 resolution area marked by a green rectangle. (c) Sketch of the Hellenic subduction zone in the wider East Mediterranean context with GPS-derived velocities with respect to Eurasia [START_REF] Kahle | GPS-derived strain rate field within the boundary zones of the Eurasian, African, and Arabian plates[END_REF]. Black barbed lines with triangles pointing towards opposite directions represent the external limit of the Mediterranean Ridge (MR) over the Ionian oceanic basin and the outer limit of the backstop (BS) from SW to NE respectively.

Geodynamical background and previous studies

The Hellenic subduction zone is marked in the southwestern part by subduction of the oldest oceanic lithosphere of the Mediterranean/Ionian Sea, at the edge of the African plate and the southwestwards fast advancing of the Aegean continental domain (~4cm/yr) [START_REF] Kahle | GPS-derived strain rate field within the boundary zones of the Eurasian, African, and Arabian plates[END_REF][START_REF] Nocquet | Present Day kinematics of the Mediterranean: A comprehensive overview of GPS results[END_REF]) (Fig. 1c). The Africa/Europe convergence vector is highly oblique to the margin and responsible for strain partitioning at the backstop edge [START_REF] Chamot Rooke | Western Mediterranean Ridge mud belt correlates with active shear strain at the prism-backstop geological contact[END_REF].

Seismic coupling in this region has been questioned since the mid-1970s, with numerous studies concluding that the plate interface accommodates up to 80% of the convergence through aseismic slip [START_REF] North | Seismic slip rates in the Mediterranean and Middle East[END_REF]Jackson and McKenzie, 1988a,b;[START_REF] Baker | Earthquake mechanisms of the Adriatic Sea and Western Greece: geodynamic implications for the oceanic subductioncontinental collision transition[END_REF]Shaw and Jackson, 2008;[START_REF] Shaw | Earthquake mechanisms and active tectonics of the Hellenic subduction zone[END_REF][START_REF] Reilinger | Geodetic constraints on the tectonic evolution of the Aegean region and strain accumulation along the Hellenic subduction zone[END_REF][START_REF] England | Constraints from GPS measurements on the dynamics of deformation in Anatolia and the Aegean[END_REF], with earthquakes of Mw ~7 rupturing isolated locked patches (Howell et al., 2017a;[START_REF] Vernant | Geodetic evidence for low coupling on the Hellenic subduction plate interface[END_REF]. However, in the absence of offshore geodetic observations the size and location of these locked patches are poorly constrained. Furthermore, local offshore seismicity is not sufficiently constrained by land-based permanent stations to provide insight on the seismogenic patterns of the interplate thrust boundary or a robust and well constrained velocity model for the area.

Local scale earthquake studies on the island of Crete with onshore seismological stations provided constraints on the distribution of seismic activity associated with the subduction beneath southern Crete [START_REF] Meier | Seismicity of the Hellenic subduction zone in the area of western and central Crete observed by temporary local seismic networks[END_REF]. For the region offshore southwest Crete and the southern Peloponnesus however, hypocenters lack offshore constraints. A year and half before the 2008 Methoni earthquake, in May-October 2006, we conducted a pilot study over that offshore domain by deploying 5 OBS jointly with coastal stations in order to better constrain the seismicity location. That study has revealed that the recorded hypocenters along the 250-km-long segment of the southwest Hellenic subduction zone occur mostly at 10-15 km depth, in the upper plate's crust [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. This first well-constrained local seismic activity was shown to correlate with an Aegean fore-arc specific feature, the Matapan Trough (MT) which has not previously been considered as a seismically active region. This major -400 km long-linear depression, reaching the greatest water depth (5km) for the Mediterranean Sea, has long been considered as the Hellenic plate boundary trench (e.g., [START_REF] Jongsma | Bathymetry and shallow structure of the Pliny and Strabo Trenches south of the Hellenic Arc[END_REF] and is still called the Hellenic Trench. It is a 4-5 km deep forearc basin which marks the transition between the inner forearc to the NE and the backstop to the SW (Le Pichon et al., 2019, Part 2) (Fig. 1a). Despite marine seismic studies, the tectonic origin of this bathymetric feature of the Aegean plate remains controversial. Once the Mediterranean Ridge was recognized as an accretionary prism [START_REF] Le Pichon | Plate boundaries and extensional tectonics[END_REF], [START_REF] Huchon | Tectonics of the Hellenic trench: a synthesis of Seabeam and submersible observations[END_REF][START_REF] Le Pichon | Plate boundaries and extensional tectonics[END_REF] interpreted the Matapan Trough as a narrow deep forearc basin. Based on Sea-Beam data and early seismic profiles, [START_REF] Lallemant | Spatial transition from compression to extension in the Western Mediterranean Ridge accretionary complex[END_REF] proposed that the MT is dominated by the along-arc extension. Based on seismic velocities and structure, and estimated time evolution of the Hellenic subduction zone, Le Pichon and [START_REF] Le Pichon | The Mediterranean Ridge backstop and the Hellenic nappes[END_REF] proposed that the backstop consists of a pile of Hellenic nappes. [START_REF] Shaw | Earthquake mechanisms and active tectonics of the Hellenic subduction zone[END_REF] considered instead the scarp of the Matapan Trough as the trenchward edge of the backstop to the deformed accretionary prism. This scarp would be the surface expression of a reverse fault, splaying off the deeper underlying thrust interface of the subduction zone. The authors propose that this fault, rather than the seismically low coupled plate interface, would be the causative fault that hosted the 365 AD Mw=8-8.5 tsunamigenic mega-event.

The few low-angle thrust events of the instrumental period highlighting the seismogenic portion of the mega-thrust subduction fault occurred mainly arcward of the Matapan Trough [START_REF] Laigle | Variation of seismic coupling with slab detachment -upper plate structure along western Hellenic subduction[END_REF][START_REF] Shaw | Earthquake mechanisms and active tectonics of the Hellenic subduction zone[END_REF]. Even though the seismogenic interplate boundary lies offshore and is accessible by seismic reflection profiling, its imaging was hindered by its significant depth (over 15 km), the presence of the sub-surface evaporites [START_REF] Chaumillon | Deformation of the western Mediterranean Ridge: Importance of Messinian evaporitic formations[END_REF]Reston et al., 2002b) plus the stacked alpine nappes of the upper plate's crust [START_REF] Le Pichon | The Mediterranean Ridge backstop and the Hellenic nappes[END_REF].

Teleseismic receiver function (RF) imaging on a dense 2D seismic array has revealed the Moho of the subducting plate, below the onshore forearc, to exhibit along-strike steps (Sachpazi et al., 2016a).

The authors showed that the slab is segmented into subducting panels by a series of nine along-dip faults. Extending the imaging further offshore by 5 OBS, [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF] proposed that segmentation continues below the 8-10 km thick backstop to its outer limit (Fig. 1a). They further suggested that these faults control the size and location of earthquakes in the megathrust boundary, such as the Methoni earthquake. This event was recorded by a nearby dense land network (Thales Was Right "TWR" Experiment) and hence could be located with the highest resolution ever obtained in that area for an event of this order of magnitude [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. The Methoni earthquake was initiated along the offshore prolongation of the intra-slab fault F6 (Fig. 1a). Its focal mechanism shows a typical subduction thrust event on a N35°E dipping plane.

The source process, studied by teleseismic and regional broad band waveform modeling [START_REF] Roumelioti | The 14 February 2008 earthquake (M6.7) sequence offshore south Peloponnese (Greece): Source models of the three strongest events[END_REF], showed that the rupture extended unilaterally towards the SE along the ~2/3 of the F6F7 panel, in two areas of high coseismic slip (Fig. 2a). Two hours after the mainshock, a large interplate event (Mw=6.5) occurred at the southeastern border of the main rupture's slip extent. One week later, another strong event (Mw=6.2) occurred in the overriding plate directly above the Mw 6.5 hypocenter with a strike slip mechanism. Significant aftershock activity continued within the overriding plate over the next four months, but was manifested as a series of reverse faulting earthquakes (M w between 4.6 and 5.4) (Fig. 2a). Their hypocenters were located across the remaining eastern part of the F6F7 panel which has not slipped co-seismically. On the other hand, no such significant aftershock activity occurred at the western part of the panel. The pronounced effect of this strong activity on seismicity rate is visible as the large first peak in Fig. 2b. After a relatively quiet 3months period, two bursts of activity mark the 2 nd phase of earthquake clustering and seismicity rate increase which corresponds to the observations discussed in this paper.

Data and Method

We deployed a total of 33 three-component ocean bottom seismometers (OBS) in the southwestern part of the Hellenic subduction zone, during a period of over 11 months, (October 2008-September 2009) (Fig. 1a). The OBS array was complemented by 15 onshore instruments. We refer to this campaign as the "post-Methoni experiment", and to previous experiment (May 2006-October 2006) as the "pre-Methoni experiment". Both campaigns were part of a major onshore/offshore experiment ("TWR" EU project 2006-2011), which also involved the collection of teleseismic data used for receiver function studies across the Peloponnesus (Sachpazi et al., 2016a) and its adjacent offshore domain [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. Observations from more than 2000 events were manually picked and the hypocenters were initially located using a simple 3-layer velocity model [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF] and the Hypo71 code [START_REF] Lee | User manual for HYPO71PC[END_REF]. To improve locations, we constructed a minimum 1D P-wave velocity model using the VELEST code [START_REF] Kissling | Program VELEST user's guide -short introduction[END_REF][START_REF] Kissling | Initial reference models in local earthquake tomography[END_REF], which jointly inverts for the velocity model along with the respective station delays and hypocenter coordinates using a damped least squares iterative inversion scheme. The procedure was performed on a selected dataset of 1104 well-locatable events with at least 7 Pobservations and azimuthal gap <180. There are many fewer events with enough S-wave picks and small gap (146 events with at least 5 P-and 5 S-wave picks on common stations, with gap < 180) and their seismic ray coverage is very sparse over most of the study area. We regarded this S data set as too small and not adequate to allow the independent construction of an 1D velocity model for Swaves. The procedure of converging towards the final 1D P-velocity model with minimum RMS error, including a series of tests to assess the stability of the produced hypocentral solutions are analytically described in Section 1 of the supplementary material.

Our final 1D P-wave velocity model (Table 1) has a fairly constant velocity in the uppermost 28km of the overriding crust (probably mixture of mostly upper continental crustal material with some underplated material). Below that depth, the effect of the 2D dipping plate interface and the oceanic Moho becomes apparent. A more pronounced step in the velocity depth function would be expected if, across the study region, there was a Moho with some limited topography, appearing roughly sub-horizontal on a regional scale, and well sampled by the available data. On the other hand, if there is a consistently dipping Moho interface with limited 3D topography, the velocity variation from normal crustal velocities to typical mantle velocities is expected to be spread across a depth range more widely than typical for a Moho. Our final model supports the latter case for the geometries in the study region.

We further reduce the relative location uncertainties in the post-Methoni 2008-2009 seismicity by applying a double-difference relocation procedure. We incorporate the hypocenters and corrected travel-times from the results of the minimum 1D model along with differential travel-times from waveform cross-correlations (see supplementary material for a detailed description of the procedure).

Results

Hypocenter Locations of the pre-Methoni and post-Methoni seismicity

Post-Methoni seismic activity is observed along the strike of the subduction zone (WNW-ESE) for over 300 km, from west of the Peloponnesus up to the south-western coast of Crete (Fig. 3a). Most of the seismicity is localized within a 10-30 km wide band beneath the Matapan Trough (MT) region. A similar pattern was observed for the pre-Methoni period (Fig. 3b). To the SE, along the deepest part of the MT, earthquakes are distributed along both sides of the depression, with most events remaining along its southern border. Towards Crete, seismicity is more scattered and the seismic rate is comparable during the time of the two experiments. Post-Methoni seismic activity is more sustained along the panel F6F7, especially in its northwestern part (Fig. 3a Inset). This contrasts with the pre-Methoni period when the Methoni coseismic slip area is nearly devoid of seismic activity (Fig. 3b). The quiescent area includes the part of fault F6 where the mainshock initiated. Instead, the fault is very active downdip of the rupture extent, with a large number of earthquakes located along this structure. To the SW, a stripe of earthquakes striking WNW is located along the updip border of the rupture (Fig. 3a).

Integration of the existing information on the plate interface depth and geometry enables the characterization of the hypocenters with respect to the two plates. We hereafter focus on panel F6F7 which hosted the Methoni earthquake. RF imaging provides constraints on the depth of the slab Moho downdip the rupture (Sachpazi et al., 2016 a,b). However, the depth of the plate interface further updip, hence along the coseismic domain is unknown. Hypocentral locations of subduction thrust earthquakes such as the Methoni event may potentially provide a first order estimate of the depth of the interplate boundary. However in the present case reported hypocenters, are based mainly on distant observations, and spread over more than 40 km both horizontally and in depth (Fig. 2a Inset map). Spatiotemporal analysis of post-Methoni seismic activity outlines 7 groups within panel F6F7 that characterize different parts of the activated region (Fig. 4a,b). The distribution of the hypocenters jointly with their evolution in time and the part of the subduction system where these occur are analytically discussed in the Appendix A. The constrained focal mechanisms (Fig. 4c) and their P, T principal axes (Figure S1) also allow for several interesting observations and are hereafter jointly discussed.

Discussion

In this work, offshore seismicity has been constrained both horizontally and in depth by our dense local array of OBS stations, during a period between 9 and 21 months after the Methoni mainshock.

Despite its moderate magnitude in comparison to megathrust earthquakes in other subduction zones, the Methoni earthquake, is the largest interplate earthquake within the 400 km long western Hellenic subduction zone for over a half-century and was followed by post-seismic slip, as large as the coseismic slip, during the 3 year period after the mainshock. According to geodetic data modeling by Howell et al. (2017a), the afterslip area was twelve times larger than the coseismic patch, over a 100km-wide segment spanning through the two panels F5-F6 and F6-F7, centered on the Methoni rupture (Fig. 1b). The poorly constrained geodetic updip and downdip limits have been chosen to correspond to the forearc backstop trenchward edge and the 35 km depth respectively. The latter is roughly located below the Messenia and Mani capes.

Here we investigate the geometry and character of the post-Methoni seismicity within panel F6F7 (Fig. 4a,b). We discuss it with respect to the shape and extension of the rupture area as well as the widely distributed post seismic deformation.

Two clusters (#6 in cyan and #7 in black) are located at depths above the updip and downdip edges of the rupture area, respectively (Fig. 4b). Hypocenters extend through a 10 km thick layer above the interplate fault with only few events reaching shallower depths. These events have not been observed by our pre-Methoni offshore-onshore deployment, which is also the case for the other clusters discussed hereafter.

There is a complete lack of interplate activity between cluster #6 and #7 over a downdip width of about 30 km corresponding to the Methoni rupture area (Fig. 4b and Fig.A.1b Profiles 1,2). It has been observed for many of the M>8 earthquakes such as Sumatra [START_REF] Hsu | Frictional afterslip following the 2005 Nias-Simeulue earthquake, Sumatra[END_REF], Chile [START_REF] Hayes | Continuing megathrust earthquake potential in Chile after the 2014 Iquique earthquake[END_REF][START_REF] Li | Joint modeling of teleseismic and tsunami wave observations to constrain the 16 September 2015 Illapel, Chile, MW 8.3 earthquake rupture process[END_REF], and Tohoku (e.g [START_REF] Asano | Spatial distribution and focal mechanisms of aftershocks of the 2011 off the Pacific coast of Tohoku earthquake[END_REF] that large co-seismic slip areas tend to have little interplate seismicity after the mainshock rupture. The distribution of hypocenters, 9-21 months after the moderate Mw=6.8 Methoni earthquake shows a similar feature. In addition, we observe that the surrounding deep upper plate volume, right above the plate interface is similarly quiet (Fig. 4b).

There is a strong concentration of hypocenters along the slab fault F6 (Fig. 4a, clusters #1 (red) and #2 (green). Coseismic slip occurring on the megathrust fault plane which is here laterally bordered by a vertical fault zone located in the upward extent of the intraslab F6 fault, may induce earthquake swarms by micro-faulting. The present study area is located at shallower depths than the Moho of the overriding plate which is identified at ~30 km depth beneath Messiniakos gulf (Sachpazi et al., 2016a).

The overlying portion of the Aegean plate is thus largely crustal. The fact that the clusters #1 (red) and #2 (green) occur in the upper-plate crust (Fig. 4b) within a vertical plane above the intra-slab fault F6, may highlight a process that triggers activity through fluid migration in a pre-stressed region above.

Though, this activity is located downdip and not over the co-seismic area along the fault F6 due to post-seismic stress relaxation (see Fig. 5b [START_REF] Lin | Stress triggering in thrust and subduction earthquakes and stress interaction between the southern San Andreas and nearby thrust and strike-slip faults[END_REF] and/or due to lateral variation of fluids in the crustal material above the fault region. In fact, many causes might contribute to the observed spatio-temporal evolution.

The slab faults have already been shown to correlate with clustered seismicity, but this occurred at the 60-80 km slab depth within the subducting crust further downdip, beneath NE Peloponnesus and have been proposed to channel slab dehydrating fluids (Sachpazi et al., 2016a). Recent results, focusing on Greece but extended also to other subduction zones, proposed that fluids may migrate along the slab further updip and upwards, towards the overriding crust to trigger upper plate seismicity (<40km depth) [START_REF] Halpaap | Earthquakes track subduction fluids from slab source to mantle wedge sink[END_REF]. Our spatio-temporal analysis has shown that the shallower cluster along the fault F6 (cluster #1, red) precedes the deeper one (cluster #2, green) (Fig. A.1a), thus the fore-mentioned mechanism looks unlikely for our observations. Yet, it is possible that dehydration occurs locally at depth and that the deeper cluster is diagnostic of the upwards channelling of fluids.

Landwards, geodetic modelling is better constrained by the GPS stations and shows significant afterslip (Howell et al., 2017a) (Fig. 1b). This is where seismicity downdip of the coseismic rupture (group #7, cluster in black) is located right above and on top of the plate interface (Fig. 4b). This is still located in the crust-to crust contact and the interplate is not yet under the mantle of the upper plate. So we can think of these aftershocks as essentially being driven by brittle creep as proposed by [START_REF] Perfettini | Postseismic relaxation driven by brittle creep: A possible mechanism to reconcile geodetic measurements and the decay rate of aftershocks, application to the Chi-Chi earthquake, Taiwan[END_REF] for deep afterslip. Slow slip in the downdip part of the slab fault F6 may be an alternative mechanism to explain the earthquake clustering and induced deformation within the upper plate (clusters #1, #2). This would be due to the differential motion between the slab panels along this fault.

Focal mechanisms and the nature of seismic deformation with respect to the plate interface

Cluster #6 (cyan), located updip of the rupture area and just above the interplate boundary, is related with reverse and strike slip events (20-23km depth) with a P axis striking NE-SW (Fig. 4c and Fig. S1), consistent with the mainshock and major aftershocks (Fig. 2a). Cluster #5 orange, located instead beneath the rupture area and within the subducting crust (Fig. 4b) contains normal-faulting events with a T-axis trending between N-S and NE-SW (26-30km depth) (Fig. 4c and Figure S1). This type of faulting is also observed in events located at the same depths in the nearby region of panel F7F8 with a T-axis trending roughly WE (Fig. 4c). shows the plate interface with the two white vertical ticks denoting the 30-km-long downdip mainshock rupture width along the western part of the panel. The bold orange line is the slab Moho assuming an 8km thick oceanic crust [START_REF] Gesret | The oceanic nature of the African slab subducted under Peloponnesus: Thin-layer resolution from multiscale analysis of teleseismic P to S converted waves[END_REF]) (c) Focal mechanisms for events within panel F6F7 and its vicinity. Beachballs with black and gray compressive quadrants correspond to events at depths related to the upper and lower plate, respectively.

In the offshore domain between the southwest Peloponnesus and Crete, focal mechanisms of the most significant earthquakes in the last 40 years show that the regional stress field corresponds to two types of faulting. The first, reverse and low-angle thrust faulting with a NE horizontal compressive axis, has been considered to reflect the SW overriding of the Aegean domain over the Ionian oceanic crust causing also the large inter-plate subduction earthquakes [START_REF] Shaw | Earthquake mechanisms and active tectonics of the Hellenic subduction zone[END_REF][START_REF] Benetatos | Focal mechanisms of shallow and intermediate depth earthquakes along the Hellenic Arc[END_REF][START_REF] Kapetanidis | Contemporary crustal stress of the Greek region deduced from earthquake focal mechanisms[END_REF]. The second, reverse faulting within the lower plate with a trench parallel compressive axis is proposed to correspond to along-arc shortening of the subducted slab [START_REF] Shaw | Earthquake mechanisms and active tectonics of the Hellenic subduction zone[END_REF].

However, at the NW part of the subduction, in the Ionian Islands, small normal-fault earthquakes with a W-E trending T-axis have been resolved at shallow depths within the slab crust, by an early OBS and land-based network [START_REF] Sachpazi | Western Hellenic subduction and Cephalonia transform: local earthquakes and plate transport and strain[END_REF]. These events found trenchward of a change in the dip of the interplate boundary [START_REF] Hirn | A traverse front with coincident normal incidence and wide angle seismic[END_REF], were interpreted as reflecting the bending of the lower plate. These types of events have also been found along the Japan trench and similarly attributed to bending-related faulting of the subducting crust [START_REF] Gamage | Offshore double-planed shallow seismic zone in the NE Japan forearc region revealed by sP depth phases recorded by regional networks[END_REF]. In the Ionian islands, in addition to the lower plate extensional earthquakes, reverse-faulting events have been located in the upper plate above the interplate boundary, exhibiting a NE striking P-axis. The contrasting focal mechanisms on either side of the seismically imaged refelctor -at 15km depth-were evidence it was the plate interface. We suggested that lower plate bending was due to the load of the southwestwards overriding upper plate which thickens landwards.

Our findings in the present study of the southwestern part of the subduction zone are consistent with these characteristics of background seismicity of the Ionian Islands. The normal fault earthquakes are in the lower plate crust and trenchwards of the thickened Hellenic continental margin and may mark the lower plate bending to underthrust the southwestwards advancing upper plate [START_REF] Le Pichon | A new approach to the opening of the Eastern Mediterranean Sea and the origin of the Hellenic Subduction Zone. Part 2 The Hellenic Subduction[END_REF]. The plate interface may be thus defined as the boundary (at the 24-25km depth range)

between the two groups of different focal mechanisms, with those around 20-23 km depth being at the base of the overriding upper plate, consistently with the proposed Methoni mainshock's hypocenter at 23km depth.

Seismic activity within the shallow upper plate

The third group (cluster #3, blue) in the shallower part of the upper plate is located 10 km above the plate interface and above the downdip coseismic border region (Fig. 4b). It displays reverse faulting events with an along-arc P-axis (Fig. 4c and Figure S1) that differs significantly from the regional NE striking horizontal compressive axis. It is possible that this seismic swarm in the hanging wall, which occurred 13 months after the mainshock, is associated with Coulomb stress transfer above the lateral variation of rupture imposed on an obliquely orientated structure [START_REF] Vallage | Localized slip and distributed deformation in oblique settings: the example of the Denali fault system, Alaska[END_REF].

Cluster #4 (yellow) exhibits a tight geographical association with the Matapan Trough. Such relation has already been noted during the pre-Methoni period, in the wider offshore forearc, though MT associated activity was not continuing into panel F6F7 (Fig. 3b). During the post-Methoni period, MT is seismically active at the 10-15 km depths range into panel F6F7. This suggests the presence of deep active faults controlling its morphological expression at the sea bottom within the outer forearc crust, which have been seismically activated in the post-Methoni period (Fig. 3a). The focal mechanisms we could constrain for two events at 15 km depth beneath MT (panel F7F8) display reverse and strike slip faulting respectively with a NE trending P-axis (Fig. 4c and Figure S1). The extensional process that was suggested for the formation of this deep outer forearc domain [START_REF] Lallemant | Spatial transition from compression to extension in the Western Mediterranean Ridge accretionary complex[END_REF] is possibly currently replaced by oblique compression. This might result from the combined effect of the ongoing collision between the backstop and the Lybian margin south of Crete, as suggested by [START_REF] Mascle | An overview of Mediterranean Ridge collisional accretionary complex as deduced from multichannel seismic data[END_REF], and from the predicted ~2cm/yr of dextral motion at the trenchward backstop's edge as suggested by [START_REF] Chamot Rooke | Western Mediterranean Ridge mud belt correlates with active shear strain at the prism-backstop geological contact[END_REF]. 20

What drives post-Methoni seismicity

We document a seismic activation of the upper plate which appears to accommodate deformation on small faults with different orientations around and above the area of mainshock slip. We also resolve the eastern part of panel F6F7 which did not slip coseismically, to be devoid of interplate activity in the 30km downdip range (Fig. A.1b Profile 3). The Mw=6.8 main shock was followed 2 hours later by the Mw=6.5 interplate aftershock at the southeastern border of the Methoni asperity which may have continued the rupture in some extent to the SE, towards slab fault F7 [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. Unfortunately available data are insufficient to model its co-seismic slip;

hence its along-dip extent is unknown. Thus, we cannot exclude the possibility that a part of the silent plate interface of the eastern panel has experienced aseismic slip as proposed by Howell et al., 2017a.

During our pre-Methoni experiment period the whole western part of the panel was shown to be seismically quiet in comparison to the eastern one (Fig. 3b), which may also be observed in the several years back, -though less well resolved-NOA seismicity (Fig. ureS2). NOA epicentral locations following our post-Methoni experiment (2010-2019) show a progressive decrease in seismic activityalso along the wider forearc-and approach to the pre-Methoni state (Fig. S3 panel F6F7). Over the Methoni coseismic domain, this may suggest progressive relocking of the asperity and strong coupling between the plates [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. The persistence of MT related seismic activity, -outside panel F6F7-during our short-term pre-and post-Methoni experiments suggests a difference in the seismic behavior along the forearc. It is possible that the same processes we proposed to drive postseismic deformation of panel F6F7, such as dehydration-embrittlement and/or aseismic slip, are behind the continuous MT related seismic activity. Clarifying the processes at work will require seabottom geodetic and seismic monitoring.

We have shown that the clustered post-Methoni seismic activity within the outer forearc crust on the one hand and the localized slip during the Methoni earthquake along the plate interface on the other hand, accommodate complementary patterns of the seismic deformation, over the megathrust boundary. Oblique convergence structures were up to now considered to be restrained at the trenchward backstop's edge [START_REF] Chamot Rooke | Western Mediterranean Ridge mud belt correlates with active shear strain at the prism-backstop geological contact[END_REF], and the Methoni sequence documents a potentially active oblique deformation up to the Matapan Trough. Our findings, on the Methoni sequence document in detail the characteristics of this part of the SW Hellenic subduction zone with a highly deformable Aegean plate overriding a segmented lower plate.

Conclusions

The results of our long-term aftershocks study of the Methoni Mw 6.8 interplate event, acquired from recordings of both OBS and land stations, provide insights into the SW Hellenic megathrust domain where large events are uncommon, though a large historical tsunamogenic event has occurred in the past [START_REF] Papazachos | The Earthquakes of Greece[END_REF][START_REF] Stiros | Seismicity of Western Crete and the destruction of the town of Kisamos at AD 365: archaeological evidence[END_REF]. We establish spatial correlation between the regional seismicity pattern and the Matapan Trough, a major -250-km-long forearc basin within our study area, already identified by our previous studies with less dense datasets [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. We discuss spatio-temporal variations of the long-term aftershocks activity, which clustered over the Methoni co-seismic and post-seismic slip area thanks to depth constraints and focal mechanisms.

We showcase clear evidences of post-Methoni crustal seismicity occurring on mega-thrust asperities and forearc crust's faults, as well as the impact of the slab segmentation (slab fault F6) in the spatial distribution. These features did not show up before the 2008 Methoni earthquake as documented by comparison with pre-Methoni seismic studies [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF]. Though the largest crustal aftershocks account for only a small fraction of the moment release and post-seismic slip documented by GPS data (Howell et al., 2017a), our results indicate that the observed post-Methoni microseismic activity is consistent with complementary post-seismic slow deformation which spreads into the forearc region rather than on the plate interface.

distributed at depths between 10 and 20km, but is mostly concentrated at 14-15km, thus confined in the upper plate (Fig Besides the activity in panel F6F7, the rest of the study area (background seismicity) contains hypocenters spreading in a wide depth range (10-30km) (Fig. 3a) and exhibits a seismicity rate of 1-2 events/day. Beneath Matapan Trough (MT), the hypocenters are mainly located in the upper plate apart from those along panel F6F7 and its vicinity for which also the lower plate presents some activity (Fig. A.1c). Hypocentral distribution along MT from our 1-year-long experiment, shows no detectable shallow seismic activity (<10km depth) that would support the presence of a seismically active separate regional outcropping fault proposed by [START_REF] Shaw | Earthquake mechanisms and active tectonics of the Hellenic subduction zone[END_REF]. prolongation, proposed by this study, assuming an 8km thick oceanic crust [START_REF] Gesret | The oceanic nature of the African slab subducted under Peloponnesus: Thin-layer resolution from multiscale analysis of teleseismic P to S converted waves[END_REF]. The plate interface depth southeastwards of slab fault F8 is obtained from a recently acquired seismic profile west of Crete [START_REF] Hussni | Pre-Stack depth Migration imaging of the Hellenic Subduction Zone[END_REF].

Fig. 1 .

 1 Fig.1. SW Hellenic subduction region. Background bathymetric data, here and in subsequent figures

Fig. 2 .

 2 Fig. 2. Overview of the 14 February 2008 Methoni earthquake sequence and slip history. (a)

  Finally, we have taken advantage of the newly constructed 1D velocity model, including stationcorrections, where available, to relocate the pre-Methoni 2006 sequence as well, in order to discuss the results in conjunction with the relocated catalogue of the 2008-2009 experiment. The hypocentral location of the 2008 Methoni mainshock and major aftershocks are also revisited, in light of the new velocity model.

Fig. 3 .

 3 Fig. 3. Map view of seismic activity beneath the offshore forearc domain of the (a) post and (b) pre-

  [START_REF] Sachpazi | Slab segmentation controls the interplate slip motion in the SW Hellenic subduction: New insight from the 2008 Mw 6.8 Methoni interplate earthquake[END_REF], used phase-arrival data from the TWR land array, in addition to permanent stations, to determine hypocentral locations for the Methoni mainshock and its largest aftershocks by a simple 3-layer model with Vp/Vs=1.75. This resulted in a focal depth of 15km for the mainshock, which the authors proposed as the plate interface depth. We re-examine the mainshock's hypocentral location, using the 1D P-wave velocity model obtained from this study. For the construction of the minimum 1D Vp model, besides OBS, we have also employed data from land-based stations (see Fig.B.1 in the supplementary material). Many of these have available data for the 2008 mainshock. The main reason the focal depth for this earthquake is so unstable is due to the fact that there were no local stations (OBS) at that time and there are significant azimuthal gaps. Thus, the trade-off between the hypocentral and velocity model parameters results in unrealistically small RMS and uncertainty estimates. In this study, we investigate the resulting range of focal depths for the mainshock by changing the only parameters for which we have control, i.e. the velocity model and distance weighting. In contrast to the seismicity of the post-Methoni experiment which was located using only P-wave data and the minimum 1-D Vp model, in the case of the mainshock we also employ a constant Vp/Vs ratio, in order to take advantage of the available S-wave arrivals, and examine how different Vp/Vs values affect the RMS error and the resulting depth. We take all these parameters into account to estimate a possible range for the focal depth of the mainshock. The resulting focal depth for the mainshock ranges from 26.5 km to31.0 km, depending mainly on the selection of Vp/Vs and the incorporation of distant stations. We obtained lower RMS values (0.27-0.28s) when limiting the distance weighting to 50-200km. Vp/Vs=1.83 yields the smallest RMS=0.27s which is also the value supported by the[START_REF] Chatelain | Etude fine de la sismicite en zone de collision continentale au moyen d'un r´eseau de stations portables: la regionHindu-Kush Pamir[END_REF] diagram for the available 2008-2009 OBS data, resulting in a focal depth of 27 km for the mainshock. Additional tests of single-event location of the 2008 mainshock with VELEST, with or without taking into account S-wave arrivals, resulted in greater focal depths(30.5-32.0km), but the RMS was also higher (0.42-0.45s). The epicentral location of the mainshock does not vary significantly in the various tests. It is clear that these various tests do not provide a way to control quantitatively the hypocenter, however, the 1D P-wave velocity model of this study is an improvement over the crude 3-layer model that was used before (model Y in Fig. B.3a in the supplementary material). Note that all the starting models (including Model Y) follow a roughly similar profile with depth (Fig. B.3a). However, starting model Y yielded the worst results (Fig. B.2, test Y2DS3). Following the above mentioned estimates we adopt a mean focal depth of 23km for the 2008 mainshock with an uncertainty of ±5km.

Fig. 4 .

 4 Fig. 4. (a) Seismicity of panel F6F7 with relocated epicenters using the double-difference method.

  We attribute this activation to a part of the post-seismic deformation which spreads induced elastic deformation into the forearc hanging wall of the interplate fault. Howell et al., 2017a model the 20mm post seismic displacement as mainly related to widespread aseismic slip of the plate interface. It is beyond the scope of this paper to interpret GPS data but we note a change in the slope of geodetic offset (see time series of stations METH and KERY at Fig.4b of Howell et al., 2017a and inset panel in Fig.2b) at the end of 2008. Most of the ~15mm post seismic displacement occurs before the end of 2008 and the remaining ~5mm continue up to the end of 2010. The onset of slow deformation coincides in timewith the initiation of the smaller clustered activity during our post-Methoni experiment (Fig.2b). The earlier large aftershocks which also occurred in the shallow upper plate took place during the first period of faster rate deformation (Fig.2b). Their focal mechanisms, strike slip as well as reverse faulting, are compatible with upper plate deformation, also shown in the kinematics of the co-seismic slip, which is driven by the SW movement of the overriding plate.

  . A.1b Profiles 2 and 3). Its evolution in time exhibits an almost constant rate of 0.3-0.4 events/day with the exception of a small burst (10-17 April) including three M L =3.5 events(Fig. A.1a).Group #5 orangeSpatial group #5 (orange) covers a 35km x 15km region of diffuse seismicity constrained at depths of 25-30km (Fig. A.1b profiles 2,3), which corresponds to the lower plate oceanic crust. It presents a very low seismicity rate of 0.1-0.2 events/day throughout the study period (Fig. A.1a). Group #6 cyan Spatial group #6 (cyan) is located updip of the Methoni rupture, at the western part of the panel (Fig.4a). Though dispersed, it clearly follows the edges of the updip extent of the coseismic slip area, with focal depths mainly in the range 17-23km, mostly concentrated at 20km depth (Fig. A.1b profiles 1,2 at distances 40-50km along the X axis) hence right above the interplate boundary. Despite its low background rate, it also contains the strongest event of the study period, an M L =4.4 earthquake that occurred on 21 April 2009 and caused a short outburst of ~20 events during the following 10 days, including an M L =4.0 event on 27 April 2009 (Fig. A.1a). This episode took place half-way through the first period of activity at cluster #3.Group #7 blackLastly, spatial group #7 (black) is located further NE, across-panel F6F7 in a NW-SE direction reaching slab fault F6 at about the same place as clusters #1 and #2, with its focal depths being distributed in the range. This group does not present any significant bursts of activity, but rather a constant seismicity rate of ~0.2 events/day which appears to slightly increase to 0.5 events/day after the initiation of cluster #3 in April 2009(Fig. A.1a).

Fig. A. 1

 1 Fig. A.1 Spatiotemporal analysis results and seismicity cross section along the Matapan Trough. (a)

Table 1 :

 1 Final 1D P-wave velocity model for the study region. Note the typical continental crustal velocities down to 28 km. Between that depth value and 43 km depth we obtain a velocity gradient reflecting the spatial variation of the depth of the Moho, which is dipping under the area (see text).

	Depth (km) Vp (km/s)
	-3.0	5.80
	5.0	5.90
	10.0	5.92
	15.0	6.26
	21.0	6.29
	28.0	6.93
	35.0	7.27

Acknowledgments

This research has been supported by the European Union FP6 NEST_INSIGHT program, under the project THALES WAS RIGHT. The authors thank the captain and the crew of R/V AIGEO of the Hellenic Center for Marine Research as well as the participants of both Géoazur OBS team (Yann Hello, Olivier Desprez, Alain Anglade and Audrey Galve) and the GEOMAR OBS team (Heidrun Kopp, Anke Dannowski and Anne Becel). We thank the National Observatory of Athens, NOA for providing data from permanent seismological stations. The authors thank the editor and two anonymous reviewers for their constructive evaluation.

Yagi, Y., Kikuchi, M., Nishimura, T., 2003. Co-seismic slip, post-seismic slip, and largest aftershock associated with the 1994 Sanriku-haruka-oki, Japan, earthquake. Geophys. Res. Lett. 30, 2177. https:// doi:10.1029/2003GL018189.

APPENDIX A

Spatiotemporal analysis of the 2008-2009 Methoni seismicity

To enable the spatiotemporal analysis of the Methoni earthquake sequence during the study period, we applied a spatial grouping methodology similar to the one followed for the relocation procedure to the final catalogue which is fully described in the Supplementary Material. Particular focus was placed in the clustered seismicity that has occurred within panel F6F7 and especially on slab fault F6, while the rest of activity in the other panels or further onshore is considered here as a "background seismicity". An inter-event distance matrix was constructed for the focused seismicity and Ward's (1963) linkage was applied to create hierarchical clustering. By selecting proper thresholds, the seismicity was initially divided in several spatial groups, which were then adjusted, merged and reduced to a selection of 7 groups that characterize different parts of the activated region within panel F6F7 (Fig. 4a). Post-Methoni activity at the beginning of the experiment (mid-October 2008), ~9 months after the mainshock (M) presented a relatively steady background rate of about 2-3 events per day (Fig.A.1a).

Group #1 red (M+10 months)

On 14 December, a sudden increase of activity was identified as group #1 (red). Epicentres were aligned along the trace of slab fault F6 beyond the downdip rupture area (Fig. 4a), but at much shallower depths of 12-17 km, hence in the upper plate (Fig.A.1b Profile 1). This spatiotemporal cluster began with smaller events was then followed by a larger one (M L =4.2) and continued for 11 days.

Group #2 green (M+11 months)

On 15 January 2009, a different cluster (#2; green) was activated on slab fault F6 (Fig. 4a), mostly in the overriding plate at depths of Profile 1). This cluster evolved into a series of small bursts including mainly a few M L >3.0 events, with the larger being an M L =3.5 event (Fig. A.1a).

29

Most of its activity was over by mid-February 2009. Although cluster #2 is located close to #1, both being downdip of the northwestern coseismic slip patch extent (Fig. 4a), they are well-separated in both space and time and have occurred at distinctly different depths. Both are located in the upper plate while the second one extends downward to reach the plate interface. They both present an apparently sub-vertical distribution; however, no clear direction of migration in depth with time could be identified within each individual cluster. By the end of the second cluster's occurrence, both groups #1 and #2 cover a column at depths 10-30km aligned horizontally along the fault system F6, in an almost linear 5-km-long segment oriented SW-NE (Fig. A.1b Profile 1). Although there is a partial overlap between their epicenters, cluster #1 appears to have occurred at the NW half of the activated zone while cluster #2 mainly occupies the SE half.

Group #3 blue (M+13 months)

Cluster #3 (blue) was initiated with an M L =3.5 event on 28 March 2009, followed by a few events during the next day (Fig. A.1a). A new burst of its activity was triggered by a couple of M L =3.9-4.0 events on 5 April, with a rate of 2-3 events/day for the next 25 days. A stronger outbreak began on 23 May 2009, following an M L =3.8 event, starting with a rate of 25 events/day which slowly diminished, but the cluster remained active until the end of the study period, in September 2009. Due to the comparable magnitudes between the largest events of this cluster, it can be characterized as a swarm.

Located ~16km SSE of clusters #1 and #2, and ~10km SE of the slab fault F6, it fills the geographical gap between the two main patches that ruptured during the 2008 Methoni mainshock (Fig. 4a). This group occurs within the overriding plate, mainly at depths between 10 and 15 km ( The other spatial groups within panel F6F7 do not present any significant spatiotemporal clustering, but are composed of diffuse seismicity that generally delineates different regions of particular interest.

Group #4 yellow

Spatial group #4 (yellow), covers a roughly linear region along ~50km SSE of cluster #3, reaching slab fault F7 (Fig. 4a). It is located at the borders of the higher slip area of the 2008 main rupture and relates geographically with the MT region across the whole panel (Fig. 4a). This area includes almost all the largest aftershocks of the first period before our post-Methoni experiment (Fig. 2a). It is


Sokhna-Mery ; Ngom

C2n

Fatima Flores-Galicia

François-Damien Delapierre

Antoine Pallandre
email: antoine.pallandre@u-psud.fr

Jean Gamby A

Isabelle Le Potier

Anne-Marie Haghiri-Gosnet
email: anne-marie.haghiri@c2n.upsaclay.fr

Electropreconcentration diagrams to optimize molecular enrichment with low counter pressure in a nanofluidic device

Keywords: Nanofluidics, Electrokinetic transport, Preconcentration diagrams Nanofluidics, Electrokinetic transport, Preconcentration diagrams

Electropreconcentration diagrams to optimize molecular enrichment with low counter pressure in a nanofluidic device.

Introduction

In the field of highly sensitive biochemical analysis, Lab-on-a-Chip (LoC) devices have to face the challenge of maintaining a good level of sensitivity and selectivity, while decreasing the amount of analyte. The analysis of highly diluted biomarkers produced at the very early stages of a disease is always facing the issue of reaching the relevant sensitivity to provide a reliable diagnosis. It is therefore crucial to couple good separation with sensitive and specific detection [1]. Separation methods that consist in controlling the spatial location of the analyte in the fluid over time [2][3][4][5][6][7] may suffer from a lack of sensitivity since there is no preconcentration. In this context developing new solutions for sample analyte preconcentration in bioanalytical fluidic devices remains a necessity. To address this prevalent issue, many focusing techniques based on electrokinetic phenomena have been proposed, among which ion concentration polarization [8][9][10][11][12][13][14][15][16][17][18][19][20], field-amplified sample stacking (FASS) [21][22][23], concentration gradient focusing [24][25][26][27] and isoelectric focusing [28][29][30]. All these techniques allow focusing analytes by exploiting the unbalanced ionic transport between anionic and cationic species due to the competition between electroosmotic flow (EOF) and electrophoretic flow (EP). This unbalanced transport under electric field is exacerbated in a nanofluidic device that integrates channels with at least one nanometric dimension [6]. In such nanofluidic device filled with a background electrolyte (BGE), this exacerbated unbalanced transport produces an enrichment-exclusion effect that is called the ion concentration polarization (CP) phenomenon [9][10]. This CP effect is mainly governed by the background electrolyte (BGE) concentration and the geometrical constraints imposed by the nanofluidic geometry. At each side of the nanochannel, CP generates BGE enrichment (cathodic reservoir) and depletion (anodic reservoir) zones with respectively a high and a low conductivity. These conductivity gradients and the corresponding shocks. These shocks are net boundaries where the analyte can locate and thus concentrate. Based on a general analytical theory of CP [9,11], the various theoretically predicted mechanisms for CP of the BGE are governed by the two key parameters:

1/ an inverse Dukhin number , which describes the ratio of bulk conductivity to surface 1 𝐷 𝑢 conductivity, defined by

, where ν i and z i are respectively the

1 𝐷 𝑢 = 𝐺 𝑏𝑢𝑙𝑘 𝐺 𝜎 = ( 𝜈 1 𝑧 1 -𝜈 2 𝑧 2 -2𝜈 1 𝜎 ) 𝐹ℎ 𝑛 𝑐 0,𝑟
mobility and the valence number of the i th ion (1 for the ion Na + or K + and 2 for the co-ion Cl -), F is the Faraday number, h n is the nanoslit height, σ is the surface charge and c 0,r is the concentration of the co-ion in the reservoir.

2/ the velocity of the co-ion (to the wall charge) nondimensionalized by the electroosmotic velocity , where ζ n is the zeta potential value assumed to be uniform along the

𝜈 * 2 = 𝜈 2 𝑧 2 𝐹η 𝜁 𝑛 𝜀
structure, η is the viscosity and ε is the permittivity.

At high values and low , values the concentrated analyte will stack at the entrance 1 𝐷 𝑢 𝜈 *

2

of the nanochannel in a regime called "CP without propagation" ( [11]. Such stacking

1 𝐷 𝑢 ≫ 1)
could not be used for real applications of preconcentration since the analyte is confined in the vicinity of the nanochannel. On the opposite, at low inverse Dukhin number ( ), the 1 𝐷 𝑢 ≪ 1 analyte will stack at the sharp interface of enrichment shock (cathodic reservoir) or focus at the interface of depletion shock (anodic reservoir) in a regime called "propagating CP". We have previously shown based on computations [12] that an anionic analyte, such as fluorescein, diluted at low BGE concentration (down to 10µM), will concentrate in the cathodic reservoir in an [9]. This type of stacking "cathodic counter gradient focusing" (CCGF) [11,16] will be named here "Cathodic Focusing" (CF) to simplify. Our simulations have also shown that a stable CF frontline can be obtained in a sharp transition region of surface charge, between propagating CF at high surface charge and stacking CS located at the nanochannel entrance at low surface charge. However, since the range of the two parameters (surface charge and electric field) to obtain this stable CF point is too narrow, only a few experiments have been reported. For example, Hluskou et al [30] have experimentally reported concentration of negatively charged BSA, with a dynamic characteristic of co-ionic species stacking on the enrichment shock, associated to a concentration factor of 100 after 200 seconds.

In order to stabilize the propagating modes at , we have previously proposed a 1 𝐷 𝑢 ≪ 1 pressure-assisted preconcentration protocol in which a hydrodynamic pressure is added to reduce (or enhance) the electroosmotic flow. This strategy has allowed a better control of the preconcentration frontline location [31]. Here, we will show that such low cathodic counterpressure allows stabilization of the CF frontline on the enrichment shock (see Fig 1 .b)). In the first part of the paper, pressure-assisted electropreconcentration of fluorescein sodium as a model anionic analyte will be presented. The role of the nanoslit length in the competition between electrophoretic velocity and the BGE velocity at the CP enrichment shock will be discussed based on experiments with different nanoslit lengths. The three different regimes of preconcentration (propagating CF -Fig. 1a), stable CF -Fig. 1b) and CS -Fig. 1c)) observed at different electric field E and counter-pressure P can be reported on "E/P" diagram. We will show how this diagram varies as a function of the nature of BGE (KCl and NaCl) at low ionic strength. Finally, to evaluate the ability of such pressure-assisted protocol to concentrate negatively charged proteins, experiments

The role of the counter-pressure on the stabilization of the frontline

Pressure-assisted electro-preconcentration consists in applying an external hydrodynamic pressure P hydro that produces an additional flow called J pressure =J P through the nanochannel. In addition to electroosmotic (J EOF ) and electrophoretic (J EP ) flows that are competing through the CP effect, this additional hydrodynamic flow (J P ) enables to shift the location of the analyte focal point in a direction that depends on the direction of the applied pressure [31]. The total flow is given by J = J EOF + J EP + J P where the additional hydrodynamic flow J P can take positive or negative values.

In the absence of a hydrodynamic pressure (J P =0), we will use the notations "classical" or "conventional" to describe electropreconcentration.

Let's first describe classical electropreconcentration with a model BGE such as KCl and NaCl. At low ionic strength C BGE =10µM and , a propagating frontline should be observed

𝜈 * 2 < 1
in the cathodic reservoir (Fig. 1a). Zangle's theory allows predicting the location of the preconcentration frontline by comparing the nondimensionalized velocity of the analyte to the 𝜈 * 𝑖 critical values for each region (nanochannel, enrichment zone and cathodic reservoir) [11]. 𝜈 * 𝑐𝑟𝑖𝑡 𝑖 When CP propagates, analytes can move inwards on either side of the enrichment shock if their mobility is low enough to travel in the direction of the bulk flow in the enrichment zone but high enough to travel against bulk flow in the reservoir following the condition:

. 𝜈 * 𝑐𝑟𝑖𝑡 𝑖,𝑒 [11].

We describe now how pressure-assisted electropreconcentration will modify the propagating frontline. In the case of an additional pressure applied from the cathode to the anode (named "cathodic pressure"), such cathodic counter-pressure reduces the role of EOF flow by promoting the electrophoretic component. Increasing J P at moderate values of pressure allows stabilization of the CF profile as function of time as described in Figure 1b). Such stable CF electropreconcentration regime is usable for applications and can produce high concentration rate whose value increases with time. At higher counter-pressure, the analyte is pushed towards the entrance of the nanochannel in a cathodic stacking CS regime (Fig. 1c)).

Materials and methods

Glass chip fabrication including glass microstructuration and bonding has been previously

extensively described by A-C. Louër et al. [31]. As shown in Figure 3a) for conventional cathodic preconcentration without any hydrodynamic counter-pressure (J P =0), a stable CF profile is observed at low field (10V/cm), whereas an expected propagating CF is observed at higher fields 20V/cm and 50V/cm. One should note that the higher the field E is, the more propagating is the profile inside the cathodic reservoir.

These experimental observations agree with the analytical calculations of Santiago [9][10] and

Zangle [11], as well as with our previous simulations [12] at very low ionic strength, which both predict the cathodic focusing (CF) profile to be unstable over time especially for high electric fields and high surface charge. and3c). For example, at 20V/cm and P=0.2 bar (Fig. 3b)), the maximum of fluorescence intensity occurring at a distance of 600 µm from the nanoslit entrance remains quite constant as function of time even after 3 minutes. Similarly, at 50V/cm and P= 0.4 bar (Fig. 3c)), after 60 seconds, the CF peak stabilizes at a larger distance from the nanoslit, namely 950µm. Also, the concentration does not decrease for longer times which reflects a stable CF regime. This result shows that an additional hydrostatic pressure can be used to tune the buffer CP effect. This influence of additional flows on CP effect was also commented by Wang et al. [32] who has observed a propagating CP in their structures, while a stable CP region could be obtained when adding a transverse EOF flux in their "H shape" device. At low electric field, adding a high counter-pressure seems to constrain the preconcentration front toward the nanoslit in the same direction as the additional hydrostatic flow. At 10V/cm and P = 0.2 bar (Fig. 3b)), or 20V/cm and P = 0.4 bar (Fig. 3c)), the profile is stacked at the entrance of the nanoslit producing a CS regime. Similar experiments have been performed with chips that integrate shorter nanoslit (L = 100 µm -results not shown here -see Figure 6 in the following section). To conclude, adding a hydrodynamic counter-pressure from the cathodic side during fluorescein preconcentration allows manipulation and stabilization in space of the preconcentration frontline.

Time dependence of the localization of the preconcentration frontline:

Analyzing the evolution as function of time of both the peak location d max and its corresponding concentration C max appears also of great interest to determine the concentration rate C max /C 0 and thus check the efficiency of the nanodevice to concentrate the analyte. Also, it allows checking the stability of the CF regime for intermediate values of counter-pressure. Figure 4 shows such an analysis of the experiments previously reported in Figure 3, for a nanoslit length L=500µm. As shown in Figure 4a) for E=20V/cm, adding a counter-pressure P in the range 0.1-0.2 bar allows stabilization of the CF regime with a fixed localization at time higher than 200 seconds.

One can observed a shift of the peak location from around 900 µm to 600 µm when P increases from 0.1 to 0.2 bars. The optimal counter-pressure P opt = 0.2 bars produces the most stable CF regime with a concentration that increases with time to reach C max = 450 µM after 10 minutes (C max /C 0 = 45). At higher pressure (P= 0.4 bars) the profile is stacked at the entrance of the nanoslit producing a CS regime as observed in Fig4.b) (dark curve). As shown in figure 4a), the higher the electric field is, the higher the optimal ΔP opt is too with P opt = 0.1 bars for E=10V/cm, P opt = 0.2 bars for E=20V/cm and P opt = 0.4 bars for E=50V/cm. However, since both fields 20V/cm and 50V/cm produce the same concentration rate C max /C 0 of about 45, the best conditions of preconcentration using KCl as BGE are 20V/cm and P opt = 0.2 bars if one wants to minimize both E and P. However, it appears that the time transition to reach stabilization of the CF regime is shorter at 50V/cm as shown by the dark curves of figures 4a) and b) corresponding to 0.4 bar. The highest the field is, the higher P has to be fixed to stabilize the preconcentration frontline. The concentration rate reaches a limit as function of field since the enrichment shock is a weak shock due to relatively low electric fields across the interface between enrichment region and cathodic reservoir [11]. ). Even if potassium can also exchange with sodium counter-ion in the KCl BGE, NaCl electrolyte appears here as a better BGE for stabilization of CF regime at low voltage. Looking at 10V/cm field and comparing left part of figure 4 and figure 5b), it appears that the NaCl BGE leads to stable CF until 0.15 bar associated to a higher C max /C 0 preconcentration factor compared the KCl BGE. We assume that the sodium cations provide rapid exchanges between the NaCl BGE and the fluorescein. This situation offers a narrower and more fluorescent CF stable window in the preconcentration diagram. With KCl BGE, the bigger and slower potassium counter-ion will slightly change the polarisation of the fluorescein and it probably creates two distinct populations of fluorescein the first majority one with Na + and the other one with K + . This configuration leads to broader but less fluorescent preconcentration plug during the CF because of the two distinct electrophoretic mobilities coming from K + -fluorescein and from the Na + -fluorescein.

It

"Electric field/Pressure" diagrams:

From all these experimental results at a unique ionic strength of 10µM, it is of great interest to build "electric field / pressure" diagrams, which were obtained in the following way. For each experiment, a classification of the different types (propagating CF, or stable CF, or CS) of preconcentration regimes obtained can be done. For each profile obtained, a point corresponding to one of the three regimes is located in an "electric field / pressure" diagram that will be named E/P diagram in the following. These E/P diagrams allow observing the three regions that correspond to the three different preconcentration regimes, namely propagating CF (in blue), stable CF (in green) and CS (in red). Four E/P diagrams have been obtained for two different lengths of the nanoslit L nanoslit = 500 µm (see Figure 6.a)) and L nanoslit = 100 µm (see Figure 6.b)) in the two BGE solutions, KCl and NaCl.

In these diagrams, the maximal value of the preconcentration rate =C max /C 0 has also been reported allowing a direct observation of the best couple (E, P) of parameters. Whatever the nature of the buffer, it is observed that the short nanoslit L = 100 μm always leads to propagating regimes when no counter-pressure is added (P=0 for conventional electropreconcentration). On the other hand, with long nanoslit (L = 500 μm), a stable CF regime can be obtained at P = 0 and low field, namely 10V / cm for KCl and 5V / cm for NaCl. This is explained by a more pronounced CP effect for 500μm-long nanoslit. Moreover, the comparison between the diagrams established for KCl and for NaCl shows that the area of observation of the stable CF regime (in green) is shifted towards the weak electric fields for NaCl. Preconcentration rates obtained are also greater with NaCl than with KCl, as shown by the values circled in red on the 4 diagrams. With NaCl as background solution, it is therefore easy to stabilize CF regime at low electric field (< or = at 5V / cm) and low counter-pressure (< or = at 0.15 bar). These conditions produce the maximal value of the preconcentration rate =C max /C 0 ~ 140. The analyst will thus have to choose between a broader stability window or a high preconcentration factor. This choice will be driven by the instrumentation, namely the speed of the camera or the sensitivity of the optical sensor. The initial concentration of the analyte is also another main parameter to consider before choosing the nature of the BGE. For trace analysis, a fluorescence optical sensor with high spatial resolution will be probably preferred since high preconcentration rate is required. In that case, NaCl appears as the best BGE. On the other hand, with analyte at higher concentration, with KCl as BGE, the stability of the preconcentration plug could be reached faster because of the wider stability domain (see Figure 6). Once again, these results underline the interdependence between the choice of the BGE and the quality of the analysis by electrophoretic methods. In other words, similarly as the popular free zone electrophoresis, electropreconcentration with a microfluidic device bearing a nanoslit requires cautious adjustments of the BGE composition regarding to the sample composition and its available quantity. Chicken ovalbumin is a 45 kDa protein that is the major component of egg white. Fluorescent ovalbumin can be used as endocytic tracers [33][34][35] and to estimate the size of the transport pathways of micro-vessels [36]. It is also used in various research fields such as studying the structure and functionality of serpins, and proteomics (egg yolk ovalbumin is frequently used as a molecular marker for calibration of electrophoresis gels) or immunology frequently used to stimulate an allergic reaction on test subjects [37][38]. For these experiments, only 500µm-long nanoslit were used as they produce the highest CP effect of the BGE in the MNM structure. Here, the background solution chosen was a real buffer frequently used in analytical chemistry, the organic zwitterionic HEPES compound. Unlike phosphate and bicarbonate buffers, this HEPES buffer solution helps maintaining the structure and function of proteins at low temperatures. Moreover, the use of zwitterionic buffers as BGE allow the application of high separation voltages while maintaining a low current. Ovalbumin was thus diluted at a concentration of 10 μM in a 10 μM HEPES solution. Ovalbumin was marked with the Texas Red marker whose maximum emission and absorption wavelengths are respectively =615nm and 596nm.

Similarly, just as sodium fluorescein, ovalbumin concentrates in the cathodic reservoir. For clarity, only profiles corresponding to a stable CF regime have been reported here in Figure 7a).

Despite the high CP effect in HEPES due to the long nanoslit (L = 500μm), propagating regimes are always observed using conventional electrophoresis at P = 0. Quite high counter-pressure must be applied, from 0.15 bar (with E = 5V / cm) to 0.4 bar (with E = 20V / cm) to stabilize the CF regime. On the other hand, similarly as for our previous experiments with NaCl as background salt, it is not necessary to apply high electric fields, as shown in the diagram obtained (Figure 7.b).

Obtaining the stable CF regime for a protein of interest such as ovalbumin at low-electric field is very interesting in terms of applications. The maximal concentration rate for ovalbumin of about  = C max /C 0 = 73 after several minutes is also promising. All those experiments carried out with ovalbumin confirm the great interest of adding an additional counter-pressure to obtain quickly

Conclusion

In this work, a pressure-assisted preconcentration protocol in which a hydrodynamic pressure is added to reduce the electroosmotic flow is shown to stabilize propagating modes at . Low cathodic counter-pressure allows stabilization of the CF frontline on the 1 𝐷 𝑢 ≪ 1 enrichment shock whereas higher counter-pressure produces CS stacking regime at the nanoslit entrance. The three different preconcentration regimes (propagating CF, stable CF and CS) observed at different electric field E and counter-pressure P for fluorescein sodium (C=10µM) can be reported on an original "E/P" diagram. The role of the nanoslit length in the competition between electrophoretic velocity and the BGE velocity at the CP enrichment shock can be easily evidenced on the corresponding "E/P" diagram. It is shown that higher P values are needed for CF stabilization with shorter nanoslit (L nanoslit = 100 µm) with a weaker CP effect, compared to longer slit (L nanoslit = 500 µm). The stable CF region in such "E/P" diagram is also changing as function of the nature of BGE (KCl and NaCl) with an easier stabilization at low E for NaCl and higher preconcentration rates. Finally, the ability of such pressure-assisted protocol to concentrate negatively charged proteins has been tested with 10µM ovalbumin in HEPES and the corresponding E/P diagram for ovalbumin confirms the existence of a stable CF regime at low electric field.

  local electromigration velocity of analytes produce a stacking/focusing phenomenon at the interfaces of enrichment (cathodic reservoir) and depletion (anodic reservoir)

  regime. Such a very low ionic strength with high surface charge promotes a cathodic concentration profile unstable over time that corresponds to the propagating CP of the BGE (see Fig.1(a))

Figure 1 . 6 CF

 16 Figure 1. Cross-view of the micro/nano/micro (MNM) device with (a) a typical "propagating CF" profile

  ovalbumin in HEPES. The E/P diagram for ovalbumin confirms the existence of a large range of parameters where a stable CF regime is obtained.

9 Figure 2

 92 Figure 2. a) schematic of the PMMA chip holder allowing mounting the nanofluidic chip on the inverted microscope and picture of the holder recorded during experiment, b) schematic of the H chip with an integrated nanochannel and top-view photo of one nanofluidic glass chip.

11 Figure 3 .

 113 Figure 3. Electropreconcentration profiles of fluorescein diluted in KCl recorded for a nanoslit with L=500µm at several additional hydrodynamic pressures (a) P=0 (conventional preconcentration), (b) P = 0.2 bars and (c) P =0.4 bars.

  counter-pressure P allows stabilizing the fluorescein concentration profile as it can be observed in Figures 3b)

Figure 4 .

 4 Figure 4. a) Evolution with time of C max the maximum concentration of each profile for 10V/cm, 20V/cm and 50V/cm given in Figure 3, and b) corresponding time dependence of its location (distance from the

Figure 5 .

 5 Figure 5. Evolution with time of C max the maximum concentration of each profile for (a) 10V/cm, (b) 20V/cm and (c) 50V/cm for experiments performed in NaCl (L=500µm) and time dependence of its location

Figure 5

 5 Figure5shows how both the peak location d max and the maximum in concentration C max

Figure 6 .

 6 Figure 6. The "Electric field /Pressure" diagram established from Fluorescein electropreconcentration

Figure 7 .

 7 Figure 7. (a) Electropreconcentration profiles of ovalbumin diluted at 10µM in HEPES at 10µM

  over time inside a micro/nano/micro fluidic device able to concentrate such derivatized biomacromolecule at high concentration rate.

Figure 1 .

 1 Figure 1. Cross-view of the micro/nano/micro (MNM) device with (a) a typical "propagating CF" profile observed for an anionic analyte stacking and moving at the enrichment shock with ν 2 * <1 for conventional electropreconcentration (without any counter-pressure ΔP hydro = 0), (b) stabilization of the concentration frontline under the application of a low cathodic counter-pressure ΔP hydro producing a stable CF regime and (c) at higher cathodic counter-pressure ( for example ΔP hydro = 4 bars), the profile is pushed towards nanoslit entrance producing a stacking CS regime

Figure 2 .

 2 Figure 2. a) schematic of the PMMA chip holder allowing mounting the nanofluidic chip on the inverted microscope and picture of the holder recorded during experiment, b) schematic of the H chip with an integrated nanochannel and top-view photo of one nanofluidic glass chip.

Figure 3 .

 3 Figure 3. Electropreconcentration profiles of fluorescein diluted in KCl recorded for a nanoslit with L=500µm at several additional hydrodynamic pressures (a) ΔP=0 (conventional preconcentration), (b) ΔP = 0.2 bars and (c) ΔP =0.4 bars.

Figure 4 .

 4 Figure 4. a) Evolution with time of C max the maximum concentration of each profile for 10V/cm, 20V/cm and 50V/cm given in Figure 3, and b) corresponding time dependence of its location (distance from the nanoslit) for 10V/cm, 20V/cm and 50V/cm. A critical time of 200 s is observed to reach the stabilization.

Figure 5 .Figure 6 .

 56 Figure 5. Evolution with time of C max the maximum concentration of each profile for (a) 10V/cm, (b) 20V/cm and (c) 50V/cm for experiments performed in NaCl (L=500µm) and time dependence of its location (distance from the nanoslit) for (d) 10V/cm, (e) 20V/cm and (f) 50V/cm.

Figure 7 .Figure S1. 2 -

 72 Figure 7. (a) Electropreconcentration profiles of ovalbumin diluted at 10µM in HEPES at 10µM recorded for a nanoslit with L=500µm at different coupled E/P that produce stable CF regimes, (b) Evolution with time of Cmax the maximum concentration of each profile for 10V/cm and 0.2 bars, and time dependence of its location (distance from the nanoslit), (c) The "Electric field /Pressure" diagram established from ovalbumin electropreconcentration experiments showing at 20V/cm and 0.4 bars the maximal rate τ=C max /C 0 =73

  is of great interest to study how the nature of the BGE can affect these preconcentration

	profiles. Similar experiments were done with NaCl as BGE ( 1 𝐷 𝑢 = 0.18 and = 0.83 -see SI.1 𝜈 *

2

for the calculations) for which a propagating regime is also predicted. With a higher inverse

Dukhin number

for NaCl compared to KCl, on should expect an easier stabilization of the 1 𝐷 𝑢 propagating CF in this BGE.

Acknowledgments

All authors would like to thank ANR for the financial support (DIFLUSEL and NANOCODE).