Tectonic tremor: the chatter of mafic underplating beneath southern Vancouver Island?

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Geena Littel, Michael Bostock, Charles Sammis, Simon Peacock, and 1 more This is a preprint; it has not been peer reviewed by a journal. https://doi.org/ 10.21203/rs.3.rs-3909443/v1 This work is licensed under a CC BY 4.0 License Status: Posted Version 1 posted You are reading this latest preprint version Abstract Tremor is a low-amplitude seismic signal that usually temporally coincides with episodic slow fault slip at plate boundaries worldwide. Since the discovery of tremor in Cascadia, significant effort has been devoted to understanding its relationship to slow slip. However, its source mechanism has been widely debated, owing in large part to the challenge of locating sources accurately in depth. We assemble a tremor catalog of 4,851 events for a ~ 10 X 20 km 2 area on southern Vancouver Island from slow slip episodes in 2003–2005 using a cross-station detection method adapted from previous studies, which we extend to use both P- and S- waves, thereby recovering accurate depths. Tremor occurs in distinct, quasi-planar clusters in the plate boundary region at a depth near 39 km, just beneath a layer of high reflectivity and within a zone of elevated Poisson’s ratio. We interpret this tremor to represent mafic underplating, wherein shearing generates tremor and continuously erodes basaltic material of the upper few hundred meters of the oceanic crust. Comminuted basalt with an increasingly anisotropic fabric is gradually plated onto the overriding lithosphere to form the highly reflective layer. Localized areas of material transfer within the subduction zone may manifest the distinct tremor clusters. Earth and environmental sciences/Solid Earth sciences/Seismology Earth and environmental sciences/Solid Earth sciences/Tectonics Figures Figure 1 Figure 2 Figure 3 Figure 4 Introduction Since the discovery of tectonic tremor in Cascadia [ 1 ], significant effort has been devoted to understanding its relationship to episodic slow fault slip (i.e., episodic tremor and slip, ETS). Tectonic tremor (or simply, “tremor”) is a low-amplitude, seismic signal bandlimited between ~ 1–10 Hz [ 2 ] [ 3 ] that usually temporally coincides with slow slip in Cascadia. While it occurs in other subduction zones and strike-slip faults worldwide (e.g., [ 4 ] [ 5 ] [ 6 ]), it is best documented in warm subduction zones such as Cascadia and Nankai (southwest Japan). Tremor is widely regarded to comprise a superposition of individual low-frequency earthquakes (LFEs) and is commonly used to infer detailed migration patterns of slow slip (e.g., [ 7 ] [ 8 ]). The source mechanism responsible for LFEs has been widely debated. Contending hypotheses include: (1) shear slip in the plate boundary zone, with focal mechanisms consistent with thrust faulting [ 9 ] [ 10 ] [ 11 ] [ 12 ], (2) slip along multiple surfaces that are distributed across ~ 40 km in depth [ 13 ], (3) rapid fluid transients or pore pressure waves (e.g., [ 14 ] [ 15 ]), or (4) local shear instabilities in a granular channel [ 16 ]. Nonetheless, the sensitivity of tremor and LFE activity to Earth tides, and the presence of a zone of elevated Vp/Vs and depressed shear-wave velocity in the tremor source region suggest that fluids play a significant role by lowering the effective stress [ 17 , 18 , 19 , 20 , 21 ]. A primary challenge in ascertaining the source process of LFEs lies in their location in depth. Depth is difficult to determine accurately because of the low amplitudes of P-waves arising from their generation at a shallowly dipping or steeply dipping structures (in subduction zones or transform faults, respectively). Radiation patterns at subduction zones favor the observation of S-waves at nearby stations (e.g., [ 22 ]). Therefore, most previous studies (e.g., [ 2 , 23 , 13 , 24 , 8 , 25 ]) used only S-waves to determine locations, frequently assuming a slab model to fix locations in depth. Where station distribution and signal levels allow, some authors have managed to identify P-waves to better constrain depths. P-waves are occasionally visible on well isolated LFEs (e.g., [ 9 ] in Japan; [ 10 ] in Cascadia), and accurate S-P times can also be recovered in high signal-to-noise ratio (SNR) circumstances using cross-correlation of vertical and horizontal recordings [ 10 , 26 ]. More generally, signal-to-noise ratios can be improved considerably by assembling templates from hundreds of LFE detections using iterative stacking and matched-filtering [ 12 , 27 ] to yield clear P- and S-arrivals that reveal features such as P-polarities [ 12 ] and S-wave splitting [ 28 ] but with the disadvantage that only average source properties (e.g. location and focal mechanism) are represented. Such studies have constrained LFE depths in Cascadia and Japan to depths near the inferred plate interface. However, higher precision and more systematic mapping of tremor hypocenters are required to better understand their origins. References [ 8 ] and [ 26 ] demonstrated how high precision tremor epicenters can be recovered from cross-station correlations of 4 s S-wave windows at three stations (3S). We extend their work using elements of [ 26 ] and information on wave propagation and radiation supplied through LFE templates to address this requirement. We generate two catalogs: one using 4 s windows of S-waves at four stations (4S) which yields the most detections but requires an assumed depth, and one using 4 S-wave plus 1 P-wave (4S + 1P) windows which gives precise depth control but yields fewer detections (see Supplementary Table 1). Several authors have noted associations between the occurrence of tremor and underplating in warm subduction zones [ 29 , 30 , 31 ]. However, a direct link between tremor and underplating has yet to be confirmed, largely due to challenges in obtaining precise locations necessary to identify signatures of underplating. Southern Vancouver Island provides an exceptional setting in which to address this problem because minimal crustal scattering yields comparatively clean seismic waveforms dominated by direct arrivals (e.g., Supplementary Fig. 1; [ 12 , 32 ], c.f. [ 33 ]). Moreover, stations from the temporary POLARIS array (2003–2005; [ 34 ]; Fig. 1 ) are situated near tremor sources and are sufficiently closely spaced to yield high cross-station correlations over short (4 s) windows for tremor arising in three major ETS episodes [ 8 , 26 , 35 ]. We focus on three LFE template locations near the axis of the POLARIS array, 053, 065, and 070 (as defined by [ 32 ]; see Fig. 1 ), which have not previously been analyzed in detail. Combined with detailed seismic reflection and tomographic imaging, accurate tremor hypocenters allow us to address the relationship of tectonic tremor to warm subduction and underplating. Results Tremor epicenter clusters In map view, the tremor is localized in four discrete, approximately planar clusters, separated by about 2 km, and labelled 1, 2, 3, and 4 in Figs. 2 and 3 . Each cluster exhibits a different dip orientation that we quantify using principal component analysis to define best-fit planes (Supplementary Table 1). We also fit one plane to all the events and two planes (one to the “northern patch” (clusters 1 and 2), and a second to the “southern patch” (clusters 3 and 4) (Supplementary Table 1). Division into northern and southern patches was motivated by their distinct geometries (Fig. 2 b) and their independent spatiotemporal rupture patterns (Fig. 3 c). Comparing the one, two, and four patch solutions, we found that two planes (northern and southern patches) best fit the data. The northern plane dips 11.2˚ at a 66.8˚ azimuth and the southern plane dips 0.6˚ at a 106.9˚ azimuth (see Supplementary Fig. 2). The average dip of the CSZ in this location is about 11˚ at a ~ N50˚E azimuth. The number of events in each cluster and its area are given in Supplementary Table 1. Tremor layer thickness It is evident in Fig. 2 (profile B) that tremor is distributed in bands that are up to ~ 400 m thick. The question arises: how much of this apparent thickness is due to location errors in depth and how much reflects the actual width of the shear zone? A comparison of absolute depths calculated from data at common timestamps but filtered in a relatively narrow (1.5-6 Hz) versus broad (1–8 Hz) frequency band exhibit a range of ~ 300 m (see Methods, Supplementary Fig. 3), implying that the actual maximum width of the shear-zone is probably on the order of 100m. A constraint on relative depths comes from the distribution of the distances of events from the best-fitting plane. These distributions are near normal with standard deviations of 243 m and 219 m and kurtoses of 5.3 and 4.2 for the northern and southern patches respectively (Supplementary Fig. 4). The kurtosis of a normal distribution is 3. The observation of values significantly greater than 3 implies abnormally broad tails that probably reflect some seismicity outside a narrow shear zone. Our best estimated width of the shear zone is less than or equal to 400 m, depending on the error in depth (see Supplementary Note 2). Spatiotemporal progression and moment estimates To assess spatio-temporal behavior of epicenters and determine the moment release, we used the larger 4S dataset. Most of the 15,986 detections are readily associated with one of the four clusters and therefore provide a more complete measure of tremor excitation (see Supplementary Fig. 5, Table 1). Spatio-temporal progression of the tremor hypocenters on 16 September 2005 (Fig. 3 ) clearly indicates that the north and south clusters behave independently. Similar behavior is observed for the ETS episodes in 2003 and 2004 (Supplementary Fig. 5). We estimated the scalar moment and moment magnitude from the energy for each tremor detection. As in [ 16 ], we observe a narrow normal distribution of magnitudes Mw = 1.60 ± 0.1, implying a log-normal distribution of scalar moments (Supplementary Fig. 6). Details on the moment and magnitude calculation are provided in Methods. Total moment released in each of the four clusters is given in Supplementary Table 1. Discussion Accurate depth determination of the individual LFEs that constitute tremor beneath Vancouver Island provide two constraints on their source: 1) their location relative to known subsurface structures and, 2) the width of the shear-zone that produces them. Comparison of tremor locations with other geophysical observations Constraints on subsurface structure in the study region are provided by Lithoprobe seismic reflection profiling [ 40 ], regional double-difference tomography incorporating LFE templates [ 41 ], and receiver function studies [ 42 , 43 ]. Figure 4 shows tremor hypocenters from this study overlain on a northward extrapolation of the Lithoprobe 84 − 02 seismic reflection profile. Panel A includes Vp while panel B includes Poisson’s ratio. A dipping zone of quasi-parallel reflectors, dubbed the “E-layer”,[ 40 ] was interpreted to represent underplated imbricate oceanic sediments and volcanics. However, tomographic studies constrain the P-wave velocities at the depth of the E-layer below southern Vancouver Island to be ≥ 7 km/s (Fig. 4 ; [ 41 ]) suggesting, instead, that this material is predominantly mafic [ 44 , 45 ]. All of the E-layer, tremor and LFE template locations lie within a zone of unusually high Poisson’s ratio anomaly (~ 0.28) that dips landward parallel to the slab, and is interpreted to indicate near-lithostatic fluid pressures in the tremor source region [ 42 , 44 ]. Our tremor locations indicate quasi-planar, segmented layers in the plate boundary region just below the E-layer, at a depth near 39 km and are consistent with those of nearby LFE template locations [ 41 , 44 ] projected into the profile. Our observation that the E-layer lies above the active locus of seismic deformation (Fig. 4 ) is contrary to reports from previous studies in this region [ 13 , 25 ]. Although we cannot rule out the possibility that some tremor mapped away from the four principal patches represents true detections, almost all detections are relatively tightly constrained to layers less than 400 m thick. Mafic underplating model for tremor Based on our locations, we interpret tremor beneath Vancouver Island to represent mafic underplating, wherein each LFE represents shear failure within mixed brittle-ductile deformation occurring in the top few hundreds of meters of crystalline oceanic crust (commonly referred to as layer 2A [ 46 ]) of the subducting Juan de Fuca plate. This layer is expected to be rich in free fluids (to ~ 4%; [ 46 , 47 ]) as it undergoes active prograde metamorphism at tremor depths (e.g., [ 47 , 48 ]). These fluids are produced by metamorphic reactions at lithostatic pore pressures (e.g., [ 49 ]) and promote material weakening and localized deformation. We expect that fluid distribution within the layer 2A will be heterogeneous as governed by the distribution and persistence of fault-controlled fluid pathways connecting to the seafloor prior to subduction [ 50 ], thus influencing where deformation is partly brittle versus where it is purely ductile. Simple shear induced through ongoing subduction causes the basaltic material to be continuously eroded leading to comminuted wear products with an increasingly anisotropic fabric, that are gradually plated onto the overriding lithosphere. Weaker volumes with higher water contents trapped by a highly anisotropic permeability are elongated by shear to produce the seismic reflectors that characterize the E-layer [ 51 ], which is mostly aseismic (e.g., Fig. 4 ). Localized areas where material transfer is occurring within the subduction zone may manifest the distinct tremor patches as seen in Figs. 2 and 3 . Support for this interpretation is multifold. Based on analyses of multiple exhumed accretionary complexes,[ 52 ] argued that exhumed, underplated basalt occurs as thin (≤ 300 m) layers derived exclusively from layer 2A, consistent with the model interpretation laid out above. In a study of the Arosa zone, plate boundary rocks exhumed from tremor depths,[ 53 ] noted that plate boundary slip occurs as frictional deformation in chlorite and talc schists surrounding blocks of metabasalt, consistent with our inference that tremor is hosted within layer 2A. Moreover, the remarkably coherent, coast-parallel distribution of tremor epicenters along the entire Cascadia margin seems unlikely were tremor associated exclusively with sediments given the variable sediment input along the margin. This observation provides additional support for a basaltic layer 2A origin of tremor beneath southern Vancouver Island. As previously noted by [ 30 ], this epicentral distribution also generally mirrors the high coastal topography associated with warm subduction zone settings, and is readily explained if tremor is a manifestation of crustal underplating [ 48 , 54 ]. Moreover, in the limited locations where crustal seismic profiling has been undertaken in the Cascadia forearc, notably from the Strait of Juan de Fuca through to central Vancouver Island [ 40 , 55 , 56 ] and in central Oregon [ 57 , 58 ], a highly reflective E-layer above the inferred slab has been identified. This suggests that it, like tremor, is present along the full Cascadia margin and is a key element of the subduction complex. Although [ 52 ] suggested a “peeling” of layer 2A in the transfer of metabasalt from subducting to overriding plates, the nature of tremor suggests an origin involving significant cataclasis [ 59 , 60 ]. For example, [ 16 ] observed log-normal distributions of LFE moments that that they argue can be explained via a model wherein shear failure at contacts between rigid grains jammed within a viscous channel generates tremor. In the current context, we interpret the granular and viscous elements of layer 2A to be associated with less altered tracts of metabasalt surrounded by a more intensely hydrated and overpressured matrix, respectively. The lognormal distribution of moments originates from the lognormal distribution of contact areas within jams expected as larger competent clasts are gradually broken down into smaller ones. As the clasts decrease in size, they become less prone to jamming, and we suggest that a scale-dependence set by layer 2A thickness contributes to the band-limitation (~ 1–10 Hz) of tremor [ 16 ] and the limited range of magnitudes observed in [ 16 ] and in this paper. As comminution proceeds, we expect increasing shear strain, ductile deformation, and gradual material transfer/transformation to the E-layer because of decreased density and strength imparted by the release of fluids [ 61 , 54 ]. We note that tremor and LFE template hypocenters lie on average 2–3 km below the base of the reflectivity that defines the E-layer in Fig. 5. We discount the possibility of location bias since the same velocity model[ 41 ] is used to locate hypocenters and migrate reflections. Rather, we argue that, at some point in the comminution and shearing process, a permeability anisotropy ``percolation” threshold is reached wherein fluids become segregated within horizontally contiguous “lenses” producing the pronounced reflectivity horizons and abrupt base that define the E-layer, perhaps through changes in dihedral angle as suggested by [ 46 ]. Although our estimates of slip within tremorgenic volumes based on Kostrov strain significantly exceed those previously reported for tremor within the ETS zone more generally [ 62 ], they nonetheless fall far short of the plate motion budget (~ 3mm versus ~ 4-5cm per ETS episode) indicating that ductile deformation must still dominate. It is likely then that the slow slip of ETS represents ductile shear persisting well into the lower reaches of the E-layer at steadily diminishing levels, both where tremor is well expressed and where it is not [ 63 ]. Distance-thickness calculation of the E-layer We assess the feasibility of E-layer assembly over realistic time periods using calculations modified from [ 40 ]. The E layer is roughly 100 km long and about 5 km thick beneath Vancouver Island, corresponding to a volume of 500 km 3 /km along strike. Roughly 1800 km of plate have been subducted beneath Vancouver Island over the last 40 Ma [ 40 ]. If a thickness H of the basalt layer was eroded away during this time to form the E-layer, the eroded volume is 1800*H km 3 /km along strike. If the E-layer was formed through basal accretion of layer 2A, then 1800H = 500, or H = 0.28 km, in rough agreement with the thickness of the layer 2A pillow basalts [ 64 , 65 ]. Moreover, the total relative displacement of the Juan de Fuca plate D and the thickness of the E-layer are consistent with the observed relation between displacement and thickness of crustal fault zones (Supplementary Fig. 7, [ 66 , 67 ]). Tremor as diagnostic of material transfer Finally, we note the additional occurrence of tectonic tremor beneath accretionary prisms in subduction zones such as Nankai [ 68 ], where underplating is also implied, and at major strike-slip faults (such as the Alpine Fault [ 69 ] or San Andreas Fault [ 4 ]). We suggest that the occurrence of tremor in these environments, as in the deep plate boundary of subduction zones, may be diagnostic of granular flow and/or material transfer in zones of high pore-fluid pressure (e.g., [ 70 , 71 ]). Methods Detection and Location We employ the P- and S-waveforms for LFE templates 053, 065, and 070 at 5 stations to determine a) delays for S- and P-wave arrivals corresponding to the nominal template location, computed using the alignment procedures in [ 72 ]; b) the splitting parameters that best reduce the S-wave particle motions on the two horizontal coordinates to rectilinear motion isolated to a single channel [ 73 ]; and c) the expected P-polarity for P-waveforms on the vertical component. These quantities are used to normalize 24 hour-long waveforms for the 4 S-wave stations (KELB (2003)/KLNB (2004 and 2005), PGC, SILB, SSIB) and 1 P-wave station (SNB) employed in this study (stations KELB and KLNB differ in location by ~ 40 m). PGC is designated as the time-stamp reference station (0 s delay) with template-dependent relative delays applied to the remaining 4 channels. Moreover, a single suite of station-specific splitting parameters is employed for all 3 templates to maintain the same S-waveforms (with different relative delays) for each case. Following Rubin and Armbruster, the normalized 24-hour waveforms are divided into 86396 4 s windows with 3 s overlap starting at midnight (PGC time) and cross-correlations are computed for all (4) pairs of S-stations with a common time stamp to maximum lags of +/- 0.4 s for a given template to mitigate against cycle skips (Supplementary Fig. 1). Thus, a tremor burst originating at the nominal template epicenter would register maximum correlation at 0 s lag for each pair, whereas non-zero lags characterize epicenters away from the template epicenter. The chosen +/- 0.4 s lag allows some overlap across the 3 template epicentral regions. A prospective detection is declared if 2 conditions are met ( [ 8 ]; see Supplementary Note 1) relating to thresholds on the values of a) the 4 possible 3-station delay time circuits (i.e. | t ij + t jk - ti k | ), and the mean cross correlation coefficient. In the event of a prospective detection, the first principal component waveform of the aligned S-waveforms is cross-correlated with the single-station P-waveform and a detection is declared upon meeting a second correlation coefficient threshold. The lag at maximum correlation enables computation of an S-P time (and therefore hypocentral depth) and lags are again restricted to lie between +/- 0.4 s. The computations above are performed for two different frequency bands: a narrow band of 1.5-6 Hz like that employed by [ 8 ], and a broader band of 1–8 Hz which yields fewer detections, but reduced likelihood of cycle skips due to the wider range of frequencies represented. There is therefore the possibility for up to 6 duplicate detections per time stamp (2 frequency bands for each of 3 templates). We employ several thresholds and statistics (described in detail in the Supplementary Note 1) to cull the tremor catalogue to a maximum of 1 detection per time stamp to emphasize tremor hypocentral patterns but minimize scatter arising from false detections. We do not attempt to eliminate repeated arrivals across consecutive overlapping waveforms, thereby allowing for a continuously evolving tremor wavefield. Each detection is thus characterized by 4 S times and 1 P time allowing for location with only 1 degree of redundancy (5 constraints on 4 hypocentral parameters). This approach yields high relative precision in location (dependent upon the precision of the relative delays), but an absolute accuracy that is dependent upon the validity of the velocity model given that stations are dominantly to one side of the source region. The choice of a suitable P-wave station is a compromise in SNR between epicentral distance and relative P-to-S radiation (assessable from template waveforms) and is best met for this region by station SNB [ 26 ]. We obtain initial locations using Hypoinverse [ 74 ] with a 1D velocity model based on the [ 41 ] 3D model at this location. After culling detections (described in Supplementary Note 1), we determine final hypocenters using double-difference relocation in tomoDD [ 75 ] with the 3D velocity model of [ 41 ]. We also compare the resulting relocations using SSIB as the P-wave detection station instead of SNB. The relocations preserve a very similar relative pattern, although due to inaccuracy of the velocity model, absolute locations are shifted slightly deeper by about 1.5 km. However, relocation of the same template waveforms as [ 41 ] using the method presented here shows good agreement with the [ 41 ] study. Furthermore, our final tremor catalog locations show good agreement with the [ 41 ] LFE family template locations, indicating that our method yields robust absolute locations (Supplementary Fig. 8). The final catalog represents 4,851 locations determined for a total of 13 days spread over 3 ETS episodes, as determined from the LFE catalogs of [ 32 ]. False Detections and Location Uncertainty All detections scattered away from the main patches were visually inspected for cycle-skipping in the waveform alignment or false detections. Only about 15% of the scattered detections were deemed true detections, and most of these events occurred shallower than 37 km depth. We do not observe any significant tremor patches distributed in depth or anywhere significantly above or below about 39 km depth. We therefore cull the dataset further to include only the events between 37 and 39.5 km depth. Visual inspection of many of the waveforms contributing the to the detections composing the primary layer structures indicates that most of these detections do not suffer from cycle-skipping. We estimate the location uncertainty by comparing the locations of common timestamp detections between the narrow & broad band data in 3 different templates (each gives a different location measurement, resulting in 6 different locations to compare). Epicentral uncertainty is estimated to be about 250 m, and depth uncertainty about 300 m (Supplementary Figs. 3, 9). We note that the nominal horizontal and depth errors from Hypoinverse are 1.15 km and 2.3 km, respectively. Moment and Moment Magnitude Calculations Following [ 8 ], we assign a consistent, coherent radiated energy metric for each 4 s, narrow-band S-detection window as $$E\left(t\right)= \frac{{S}_{A}\left(t\right){S}_{B}\left({t}_{B}^{{\prime }}\right)+{S}_{A}\left(t\right){S}_{C}\left({t}_{C}^{{\prime }}\right)+{S}_{B}\left({t}_{B}^{{\prime }}\right){S}_{C}\left({t}_{C}{\prime }\right)}{3}$$ , where S ( t ) is the S-wave seismogram, A , B , and C denote stations PGC, SILB, SSIB, and t ’ is the time offset between the tremor arrival at stations B or C and station A . 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Christensen, "Seismicity, Metamorphism, and Fluid Evolution Across the Northern Cascadia Fore Arc," Geochemistry, Geophysics, Geosystems, vol. 19, no. 6, pp. 1881-1897, 2018. P. Audet, M. G. Bostock, D. C. Boyarko, M. R. Brudzinski and R. M. Allen, "Slab morphology in the Cascadia fore arc and its relation to episodic tremor and slip," Journal of Geophysical Research: Solid Earth , vol. 115, no. B4, 2010. W. Bloch, M. G. Bostock and P. Audet, "A Cascadia Slab Model From Receiver Functions," Geochemistry, Geophysics, Geosystems, vol. 24, no. 10, p. e2023GC011088, 2023. A. J. Calvert, M. G. Bostock, G. Savard and M. J. Unsworth, "Cascadia low frequency earthquakees at the base of an overpressured subduction shear zone," Nature Communications, vol. 11, no. 3874, 2020. A. J. Calvert, L. A. Preston and A. M. Farahbod, "Sedimentary underplating at the Cascadia mantle-wedge corner revealed by seismic imaging," Nature Geoscience, vol. 4, pp. 545-548, 2011. R. D. 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Trehu and e. al., "Wide-Angle Seismic Recordings from the 1998 Seismic Hazards Investigation of Puget Sound (SHIPS), Western Washington and British Columbia," USGS Open-File Report , Vols. 99-314, p. 110 pp., 1999. M. R. Nedimović, R. D. Hyndman, K. Ramachandran and G. D. Spence, "Reflection signature of seismic and aseismic slip on the northern Cascadia subduction interface," Nature, vol. 424, pp. 416-420, 2003. A. M. Trehu, I. Asudeh, T. M. Brocher, J. H. Luetgert, W. D. Mooney, J. L. Nabelek and Y. Nakamura, "Crustal Architecture of the Cascadia Forearc," Science, vol. 266, no. 5183, pp. 237-243, 1994. R. W. Keach, J. E. Oliver, L. D. Brown and S. Kaufman, "Cenozoic active margin and shallow Cascades structure: COCORP results from western Oregon," GSA Bulletin, vol. 101, no. 6, pp. 783-794, 1989. J. D. Kirkpatrick, Å. Fagereng and D. R. Shelly, "Geological constraints on the mechanisms of slow earthquakes," Nature Reviews Earth & Environment, vol. 2, pp. 285-301, 2021. O. Oncken, S. Angiboust and G. Dresen, "Slow slip in subduction zones: Reconciling deformation fabrics with instrumental observations and laboratory results," Geosphere, vol. 18, no. 1, pp. 104-129, 2022. T. Gerya and F. I. Meilick, "Geodynamic regimes of subduction under an active margin: effects of rheological weakening by fluids and melts," Journal of Metamorphic Geology, vol. 29, no. 1, pp. 7-31, 2011. H. Kao, K. Wang, H. Dragert, J. Y. Kao and G. Rogers, "Estimating seismic moment magnitude (Mw) of tremor bursts in northern Cascadia: Implications for the "seismic efficiency" of episodic tremor and slip," Geophysical Research Letters, vol. 37, no. 19, 2010. A. G. Wech and N. M. Bartlow, "Slip rate and tremor genesis in Cascadia," Geophysical Research Letters, vol. 41, no. 2, pp. 392-398, 2014. R. Houtz and J. Ewing, "Upper crustal structure as a function of plate age," Journal of Geophysical Research: Solid Earth, vol. 81, no. 14, pp. 2490-2498, 1976. R. E. Houtz, "Seismic properties of layer 2A in the Pacific," Journal of Geophysical Research: Solid Earth, vol. 81, no. 35, pp. 6231-6331, 1976. C. H. Scholz, "Wear and gouge formation in brittle faulting," Geology, vol. 15, pp. 493-495, 1987. J. Hull, "Thickness-displacement relationships for deformation zones," Journal of Structural Geology, vol. 10, no. 4, pp. 431-435, 1988. K. Obana and S. Kodaira, "Low-frequency tremors associated with reverse faults in a shallow accretionary prism," Earth and Planetary Science Letters, vol. 287, no. 1-2, pp. 168-174, 2009. A. G. Wech, C. M. Boese, T. A. Stern and J. Townend, "Tectonic tremor and deep slow slip on the Alpine Fault," Geophysical Research Letters, vol. 39, no. 10, 2012. D. M. Saffer and H. J. Tobin, "Hydrogeology and Mechanics of Subduction Zone Forearcs: Fluid Flow and Pore Pressure," Annual Review of Earth and Planetary Sciences, vol. 39, pp. 157-186, 2011. G. C. Beroza and S. Ide, "Slow Earthquakes and Nonvolcanic Tremor," Annual Review of Earth and Planetary Sciences, vol. 39, pp. 271-296, 2011. M. G. Bostock, A. P. Plourde, D. Drolet and G. F. Littel, "Multichannel Alignment of S-waves," Bulletin of the Seismological Society of America, vol. 112, no. 1, pp. 133-142, 2022. Y. Peng, A. M. Rubin, M. G. Bostock and J. G. Armbruster, "High-resolution imaging of rapid tremor migrations beneath southern Vancouver Island using cross-station correlations," Journal of Geophysical Research: Solid Earth, vol. 120, no. 6, pp. 4317-4332, 2015. F. W. Klein, "User's Guide to HYPOINVERSE–2000, a Fortran Program to Solve for Earthquake Locations and Magnitudes," U.S. Geological Survey Open-File Report 02-171, p. 123 pp., 2002. H. Zhang and C. H. Thurber, "Double-Difference Tomography: The Method and its Application to the Hayward Fault, California," Bulletin of the Seismological Society of America, vol. 93, no. 5, pp. 1875-1889, 2003. K. Obana, M. Scherwath, Y. Yamamoto, S. Kodaira, K. Wang, G. Spence, M. Riedel and H. Kao, "Earthquake Activity in Northern Cascadia Subduction Zone Off Vancouver Island Revealed by Ocean‐Bottom Seismograph Observations," Bulletin of the Seismological Society of America, vol. 105, pp. 489-495, 2015. M. G. Bostock, A. A. Royer, E. H. Hearn and S. M. Peacock, "Low frequency earthquakes below southern Vancouver Island," Geochemistry, Geophysics, Geosystems, vol. 13, no. 11, 2012. A. J. Calvert, M. G. Bostock, G. Savard and M. J. Unsworth, "Cascadia low frequency earthquakes at the base of an overpressured subduction shear zone," Nature Communications, vol. 11, no. 3874, 2020. Additional Declarations There is NO Competing Interest. 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Also discoverable on Platform About Our Team In Review Editorial Policies Advisory Board Help Center Resources Author Services Accessibility API Access RSS feed Manage Cookie Preferences © Research Square 2026 | ISSN 2693-5015 (online) Privacy Policy Terms of Service Do Not Sell My Personal Information {"props":{"pageProps":{"initialData":{"identity":"rs-3909443","acceptedTermsAndConditions":true,"allowDirectSubmit":true,"archivedVersions":[],"articleType":"Article","associatedPublications":[],"authors":[{"id":270050283,"identity":"6055bbac-39d0-4687-831a-fcf3c5494293","order_by":0,"name":"Geena Littel","email":"data:image/png;base64,iVBORw0KGgoAAAANSUhEUgAAAZAAAAAyAQMAAABI0h/eAAAABlBMVEX///8AAABVwtN+AAAACXBIWXMAAA7EAAAOxAGVKw4bAAAA+klEQVRIiWNgGAWjYNCCAgkgwdxwIKECRDM3EKHFQEKCh4Gx8cCHMyAtjERpYQBpaT44sw3EI6BF3r352IMfBhZ19uyNDYd559VG87cDtfyo2IZTi+GZY+mGPSCH8RwEatl2PHfGYcYGxp4zt3FrmZFjJsED0iKRCNJyLLcBqIWZsQ2Plvnvv0n+AWmRfwjUMudY7nxCWuQleNikIbYwNhyc2VCTu4GQFgOeNDNpGQMJyZ4ziQ0HPhw7kLsRqOUgPr/Itx9+Jvmmoo6fvf3w4Q8JNXW5884fPvjgRwUeWw6g8g+DyQMY6pBtaUDl1+FTPApGwSgYBSMUAABEKlv7OFMVqAAAAABJRU5ErkJggg==","orcid":"https://orcid.org/0000-0001-5432-325X","institution":"University of British Columbia","correspondingAuthor":true,"prefix":"","firstName":"Geena","middleName":"","lastName":"Littel","suffix":""},{"id":270050284,"identity":"639eae7d-d887-4d31-a074-e43d032cf559","order_by":1,"name":"Michael Bostock","email":"","orcid":"","institution":"University of British Columbia","correspondingAuthor":false,"prefix":"","firstName":"Michael","middleName":"","lastName":"Bostock","suffix":""},{"id":270050285,"identity":"868d0f7c-b66d-4265-8393-c1995bdbbfa5","order_by":2,"name":"Charles Sammis","email":"","orcid":"","institution":"University of British Columbia","correspondingAuthor":false,"prefix":"","firstName":"Charles","middleName":"","lastName":"Sammis","suffix":""},{"id":270050286,"identity":"97858934-3c64-4beb-8326-82c3436e6823","order_by":3,"name":"Simon Peacock","email":"","orcid":"","institution":"University of British Columbia","correspondingAuthor":false,"prefix":"","firstName":"Simon","middleName":"","lastName":"Peacock","suffix":""},{"id":270050287,"identity":"f68c837a-5cb6-4e0d-9108-60c2e0de1b22","order_by":4,"name":"Andrew Calvert","email":"","orcid":"https://orcid.org/0000-0002-1024-6640","institution":"Simon Fraser University","correspondingAuthor":false,"prefix":"","firstName":"Andrew","middleName":"","lastName":"Calvert","suffix":""}],"badges":[],"createdAt":"2024-01-29 21:55:57","currentVersionCode":1,"declarations":"","doi":"10.21203/rs.3.rs-3909443/v1","doiUrl":"https://doi.org/10.21203/rs.3.rs-3909443/v1","draftVersion":[],"editorialEvents":[],"editorialNote":"","failedWorkflow":false,"files":[{"id":50932210,"identity":"922e4a07-acc0-4b0e-bf58-cdb957ce80bc","added_by":"auto","created_at":"2024-02-09 18:52:53","extension":"jpeg","order_by":1,"title":"Figure 1","display":"","copyAsset":false,"role":"figure","size":2208055,"visible":true,"origin":"","legend":"\u003cp\u003eMap of the Cascadia region with the purple line denoting the outline of tremor epicenters from 2008-2019 [36] and major tectonic boundaries (black lines) and Cascade volcanoes (triangles) indicated. Map inset shows the study region. Large blue circles show the three LFE templates used in this study; smaller blue circles are other LFE template locations [37]. Red triangles are the five recording stations used and black line is the seismic reflection profile 84-02 [38]. NFZ = Nootka Fault Zone.\u003c/p\u003e","description":"","filename":"floatimage1.jpeg","url":"https://assets-eu.researchsquare.com/files/rs-3909443/v1/bb8c187357a1f2afee7eab2d.jpeg"},{"id":50932212,"identity":"4285056d-d830-4ccb-802c-640dd0c94c49","added_by":"auto","created_at":"2024-02-09 18:52:54","extension":"png","order_by":2,"title":"Figure 2","display":"","copyAsset":false,"role":"figure","size":110528,"visible":true,"origin":"","legend":"\u003cp\u003eMap view of 4S+1P tremor hypocenters with lines A, B, and C indicating the profile sections to the right. White dots with number labels indicate respective LFE template locations. Thin black contours are Moho depths of the [39] slab model. a-c) Depth sections of tremor hypocenters along lines A, B, and C. Hypocenters are colored by patch number.\u003c/p\u003e","description":"","filename":"2.png","url":"https://assets-eu.researchsquare.com/files/rs-3909443/v1/8cf9e6457177f7a325d3bf0b.png"},{"id":50932214,"identity":"82334592-a620-45a9-89cd-ef2e576d044d","added_by":"auto","created_at":"2024-02-09 18:52:54","extension":"png","order_by":3,"title":"Figure 3","display":"","copyAsset":false,"role":"figure","size":192393,"visible":true,"origin":"","legend":"\u003cp\u003e\u003cstrong\u003ea) \u003c/strong\u003eMap view of “4S” dataset epicenters and four clusters defined for slip/moment calculations. \u003cstrong\u003eb) \u003c/strong\u003eTemporal progression of tremor hypocenters (4S dataset) color-coded by time on a particularly active day of the 2005 ETS episode (16 September). Diagonal black dashed line indicates the profile used for panel \u003cstrong\u003ec)\u003c/strong\u003e, which shows the along-profile distance of locations for the same time-period (16 September 2005).\u003c/p\u003e","description":"","filename":"3.png","url":"https://assets-eu.researchsquare.com/files/rs-3909443/v1/b1695290f6c99fedbefb91f8.png"},{"id":50932213,"identity":"573385a1-3430-4495-86bd-36df232e361d","added_by":"auto","created_at":"2024-02-09 18:52:54","extension":"png","order_by":4,"title":"Figure 4","display":"","copyAsset":false,"role":"figure","size":314109,"visible":true,"origin":"","legend":"\u003cp\u003eDepth section of tremor (green dots) overlain on P-wave velocity (left) and Poissons’ ratio (right) [37] and the seismic reflection profile 84-02 [38], modified from [44]. LFE template locations are in blue circles and forearc crustal seismicity are in black circles. Grey line is a wide-angle reflector interpreted by [44] to represent the oceanic Moho. F = reflection horizon interpreted by [38] to represent the Moho.\u003c/p\u003e","description":"","filename":"4.png","url":"https://assets-eu.researchsquare.com/files/rs-3909443/v1/be4f53d86ed383bd48d330a9.png"},{"id":77155954,"identity":"963c76f0-4382-42b3-8f90-f74cf79fff4b","added_by":"auto","created_at":"2025-02-25 16:20:10","extension":"pdf","order_by":0,"title":"","display":"","copyAsset":false,"role":"manuscript-pdf","size":3445140,"visible":true,"origin":"","legend":"","description":"","filename":"manuscript.pdf","url":"https://assets-eu.researchsquare.com/files/rs-3909443/v1/38cd6024-88f9-4055-a70d-6a0906803ec1.pdf"},{"id":50932211,"identity":"5d8bc24f-f4bf-4b7a-af7b-d20ba2759f31","added_by":"auto","created_at":"2024-02-09 18:52:54","extension":"docx","order_by":1,"title":"","display":"","copyAsset":false,"role":"supplement","size":4359802,"visible":true,"origin":"","legend":"\u003cp\u003eSupplementary Materials\u003c/p\u003e","description":"","filename":"Tremor2024SuppNote7GL.docx","url":"https://assets-eu.researchsquare.com/files/rs-3909443/v1/4e4dd451bcda6017d881991a.docx"}],"financialInterests":"There is \u003cb\u003eNO\u003c/b\u003e Competing Interest.","formattedTitle":"Tectonic tremor: the chatter of mafic underplating beneath southern Vancouver Island?","fulltext":[{"header":"Introduction","content":"\u003cp\u003eSince the discovery of tectonic tremor in Cascadia [\u003cspan citationid=\"CR1\" class=\"CitationRef\"\u003e1\u003c/span\u003e], significant effort has been devoted to understanding its relationship to episodic slow fault slip (i.e., episodic tremor and slip, ETS). Tectonic tremor (or simply, \u0026ldquo;tremor\u0026rdquo;) is a low-amplitude, seismic signal bandlimited between ~\u0026thinsp;1\u0026ndash;10 Hz [\u003cspan citationid=\"CR2\" class=\"CitationRef\"\u003e2\u003c/span\u003e] [\u003cspan citationid=\"CR3\" class=\"CitationRef\"\u003e3\u003c/span\u003e] that usually temporally coincides with slow slip in Cascadia. While it occurs in other subduction zones and strike-slip faults worldwide (e.g., [\u003cspan citationid=\"CR4\" class=\"CitationRef\"\u003e4\u003c/span\u003e] [\u003cspan citationid=\"CR5\" class=\"CitationRef\"\u003e5\u003c/span\u003e] [\u003cspan citationid=\"CR6\" class=\"CitationRef\"\u003e6\u003c/span\u003e]), it is best documented in warm subduction zones such as Cascadia and Nankai (southwest Japan). Tremor is widely regarded to comprise a superposition of individual low-frequency earthquakes (LFEs) and is commonly used to infer detailed migration patterns of slow slip (e.g., [\u003cspan citationid=\"CR7\" class=\"CitationRef\"\u003e7\u003c/span\u003e] [\u003cspan citationid=\"CR8\" class=\"CitationRef\"\u003e8\u003c/span\u003e]).\u003c/p\u003e \u003cp\u003eThe source mechanism responsible for LFEs has been widely debated. Contending hypotheses include: (1) shear slip in the plate boundary zone, with focal mechanisms consistent with thrust faulting [\u003cspan citationid=\"CR9\" class=\"CitationRef\"\u003e9\u003c/span\u003e] [\u003cspan citationid=\"CR10\" class=\"CitationRef\"\u003e10\u003c/span\u003e] [\u003cspan citationid=\"CR11\" class=\"CitationRef\"\u003e11\u003c/span\u003e] [\u003cspan citationid=\"CR12\" class=\"CitationRef\"\u003e12\u003c/span\u003e], (2) slip along multiple surfaces that are distributed across ~\u0026thinsp;40 km in depth [\u003cspan citationid=\"CR13\" class=\"CitationRef\"\u003e13\u003c/span\u003e], (3) rapid fluid transients or pore pressure waves (e.g., [\u003cspan citationid=\"CR14\" class=\"CitationRef\"\u003e14\u003c/span\u003e] [\u003cspan citationid=\"CR15\" class=\"CitationRef\"\u003e15\u003c/span\u003e]), or (4) local shear instabilities in a granular channel [\u003cspan citationid=\"CR16\" class=\"CitationRef\"\u003e16\u003c/span\u003e]. Nonetheless, the sensitivity of tremor and LFE activity to Earth tides, and the presence of a zone of elevated Vp/Vs and depressed shear-wave velocity in the tremor source region suggest that fluids play a significant role by lowering the effective stress [\u003cspan citationid=\"CR17\" class=\"CitationRef\"\u003e17\u003c/span\u003e, \u003cspan citationid=\"CR18\" class=\"CitationRef\"\u003e18\u003c/span\u003e, \u003cspan citationid=\"CR19\" class=\"CitationRef\"\u003e19\u003c/span\u003e, \u003cspan citationid=\"CR20\" class=\"CitationRef\"\u003e20\u003c/span\u003e, \u003cspan citationid=\"CR21\" class=\"CitationRef\"\u003e21\u003c/span\u003e].\u003c/p\u003e \u003cp\u003eA primary challenge in ascertaining the source process of LFEs lies in their location in depth. Depth is difficult to determine accurately because of the low amplitudes of P-waves arising from their generation at a shallowly dipping or steeply dipping structures (in subduction zones or transform faults, respectively). Radiation patterns at subduction zones favor the observation of S-waves at nearby stations (e.g., [\u003cspan citationid=\"CR22\" class=\"CitationRef\"\u003e22\u003c/span\u003e]). Therefore, most previous studies (e.g., [\u003cspan citationid=\"CR2\" class=\"CitationRef\"\u003e2\u003c/span\u003e, \u003cspan citationid=\"CR23\" class=\"CitationRef\"\u003e23\u003c/span\u003e, \u003cspan citationid=\"CR13\" class=\"CitationRef\"\u003e13\u003c/span\u003e, \u003cspan citationid=\"CR24\" class=\"CitationRef\"\u003e24\u003c/span\u003e, \u003cspan citationid=\"CR8\" class=\"CitationRef\"\u003e8\u003c/span\u003e, \u003cspan citationid=\"CR25\" class=\"CitationRef\"\u003e25\u003c/span\u003e]) used only S-waves to determine locations, frequently assuming a slab model to fix locations in depth.\u003c/p\u003e \u003cp\u003eWhere station distribution and signal levels allow, some authors have managed to identify P-waves to better constrain depths. P-waves are occasionally visible on well isolated LFEs (e.g., [\u003cspan citationid=\"CR9\" class=\"CitationRef\"\u003e9\u003c/span\u003e] in Japan; [\u003cspan citationid=\"CR10\" class=\"CitationRef\"\u003e10\u003c/span\u003e] in Cascadia), and accurate S-P times can also be recovered in high signal-to-noise ratio (SNR) circumstances using cross-correlation of vertical and horizontal recordings [\u003cspan citationid=\"CR10\" class=\"CitationRef\"\u003e10\u003c/span\u003e, \u003cspan citationid=\"CR26\" class=\"CitationRef\"\u003e26\u003c/span\u003e]. More generally, signal-to-noise ratios can be improved considerably by assembling templates from hundreds of LFE detections using iterative stacking and matched-filtering [\u003cspan citationid=\"CR12\" class=\"CitationRef\"\u003e12\u003c/span\u003e, \u003cspan citationid=\"CR27\" class=\"CitationRef\"\u003e27\u003c/span\u003e] to yield clear P- and S-arrivals that reveal features such as P-polarities [\u003cspan citationid=\"CR12\" class=\"CitationRef\"\u003e12\u003c/span\u003e] and S-wave splitting [\u003cspan citationid=\"CR28\" class=\"CitationRef\"\u003e28\u003c/span\u003e] but with the disadvantage that only average source properties (e.g. location and focal mechanism) are represented. Such studies have constrained LFE depths in Cascadia and Japan to depths near the inferred plate interface. However, higher precision and more systematic mapping of tremor hypocenters are required to better understand their origins. References [\u003cspan citationid=\"CR8\" class=\"CitationRef\"\u003e8\u003c/span\u003e] and [\u003cspan citationid=\"CR26\" class=\"CitationRef\"\u003e26\u003c/span\u003e] demonstrated how high precision tremor epicenters can be recovered from cross-station correlations of 4 s S-wave windows at three stations (3S). We extend their work using elements of [\u003cspan citationid=\"CR26\" class=\"CitationRef\"\u003e26\u003c/span\u003e] and information on wave propagation and radiation supplied through LFE templates to address this requirement. We generate two catalogs: one using 4 s windows of S-waves at four stations (4S) which yields the most detections but requires an assumed depth, and one using 4 S-wave plus 1 P-wave (4S\u0026thinsp;+\u0026thinsp;1P) windows which gives precise depth control but yields fewer detections (see Supplementary Table\u0026nbsp;1).\u003c/p\u003e \u003cp\u003eSeveral authors have noted associations between the occurrence of tremor and underplating in warm subduction zones [\u003cspan citationid=\"CR29\" class=\"CitationRef\"\u003e29\u003c/span\u003e, \u003cspan citationid=\"CR30\" class=\"CitationRef\"\u003e30\u003c/span\u003e, \u003cspan citationid=\"CR31\" class=\"CitationRef\"\u003e31\u003c/span\u003e]. However, a direct link between tremor and underplating has yet to be confirmed, largely due to challenges in obtaining precise locations necessary to identify signatures of underplating. Southern Vancouver Island provides an exceptional setting in which to address this problem because minimal crustal scattering yields comparatively clean seismic waveforms dominated by direct arrivals (e.g., Supplementary Fig.\u0026nbsp;1; [\u003cspan citationid=\"CR12\" class=\"CitationRef\"\u003e12\u003c/span\u003e, \u003cspan citationid=\"CR32\" class=\"CitationRef\"\u003e32\u003c/span\u003e], c.f. [\u003cspan citationid=\"CR33\" class=\"CitationRef\"\u003e33\u003c/span\u003e]). Moreover, stations from the temporary POLARIS array (2003\u0026ndash;2005; [\u003cspan citationid=\"CR34\" class=\"CitationRef\"\u003e34\u003c/span\u003e]; Fig.\u0026nbsp;\u003cspan refid=\"Fig1\" class=\"InternalRef\"\u003e1\u003c/span\u003e) are situated near tremor sources and are sufficiently closely spaced to yield high cross-station correlations over short (4 s) windows for tremor arising in three major ETS episodes [\u003cspan citationid=\"CR8\" class=\"CitationRef\"\u003e8\u003c/span\u003e, \u003cspan citationid=\"CR26\" class=\"CitationRef\"\u003e26\u003c/span\u003e, \u003cspan citationid=\"CR35\" class=\"CitationRef\"\u003e35\u003c/span\u003e]. We focus on three LFE template locations near the axis of the POLARIS array, 053, 065, and 070 (as defined by [\u003cspan citationid=\"CR32\" class=\"CitationRef\"\u003e32\u003c/span\u003e]; see Fig.\u0026nbsp;\u003cspan refid=\"Fig1\" class=\"InternalRef\"\u003e1\u003c/span\u003e), which have not previously been analyzed in detail. Combined with detailed seismic reflection and tomographic imaging, accurate tremor hypocenters allow us to address the relationship of tectonic tremor to warm subduction and underplating.\u003c/p\u003e \u003cp\u003e \u003c/p\u003e"},{"header":"Results","content":"\u003cdiv id=\"Sec3\" class=\"Section2\"\u003e\n\u003ch2\u003eTremor epicenter clusters\u003c/h2\u003e\n\u003cp\u003eIn map view, the tremor is localized in four discrete, approximately planar clusters, separated by about 2 km, and labelled 1, 2, 3, and 4 in Figs.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e2\u003c/span\u003e and \u003cspan class=\"InternalRef\"\u003e3\u003c/span\u003e. Each cluster exhibits a different dip orientation that we quantify using principal component analysis to define best-fit planes (Supplementary Table\u0026nbsp;1). We also fit one plane to all the events and two planes (one to the \u0026ldquo;northern patch\u0026rdquo; (clusters 1 and 2), and a second to the \u0026ldquo;southern patch\u0026rdquo; (clusters 3 and 4) (Supplementary Table\u0026nbsp;1). Division into northern and southern patches was motivated by their distinct geometries (Fig.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e2\u003c/span\u003eb) and their independent spatiotemporal rupture patterns (Fig.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e3\u003c/span\u003ec). Comparing the one, two, and four patch solutions, we found that two planes (northern and southern patches) best fit the data. The northern plane dips 11.2˚ at a 66.8˚ azimuth and the southern plane dips 0.6˚ at a 106.9˚ azimuth (see Supplementary Fig.\u0026nbsp;2). The average dip of the CSZ in this location is about 11˚ at a\u0026thinsp;~\u0026thinsp;N50˚E azimuth. The number of events in each cluster and its area are given in Supplementary Table\u0026nbsp;1.\u003c/p\u003e\n\u003c/div\u003e\n\u003cdiv id=\"Sec4\" class=\"Section2\"\u003e\n\u003ch2\u003eTremor layer thickness\u003c/h2\u003e\n\u003cp\u003eIt is evident in Fig.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e2\u003c/span\u003e (profile B) that tremor is distributed in bands that are up to ~\u0026thinsp;400 m thick. The question arises: how much of this apparent thickness is due to location errors in depth and how much reflects the actual width of the shear zone? A comparison of absolute depths calculated from data at common timestamps but filtered in a relatively narrow (1.5-6 Hz) versus broad (1\u0026ndash;8 Hz) frequency band exhibit a range of ~\u0026thinsp;300 m (see Methods, Supplementary Fig.\u0026nbsp;3), implying that the actual maximum width of the shear-zone is probably on the order of 100m. A constraint on relative depths comes from the distribution of the distances of events from the best-fitting plane. These distributions are near normal with standard deviations of 243 m and 219 m and kurtoses of 5.3 and 4.2 for the northern and southern patches respectively (Supplementary Fig.\u0026nbsp;4). The kurtosis of a normal distribution is 3. The observation of values significantly greater than 3 implies abnormally broad tails that probably reflect some seismicity outside a narrow shear zone. Our best estimated width of the shear zone is less than or equal to 400 m, depending on the error in depth (see Supplementary Note 2).\u003c/p\u003e\n\u003c/div\u003e\n\u003cdiv id=\"Sec5\" class=\"Section2\"\u003e\n\u003ch2\u003eSpatiotemporal progression and moment estimates\u003c/h2\u003e\n\u003cp\u003eTo assess spatio-temporal behavior of epicenters and determine the moment release, we used the larger 4S dataset. Most of the 15,986 detections are readily associated with one of the four clusters and therefore provide a more complete measure of tremor excitation (see Supplementary Fig.\u0026nbsp;5, Table\u0026nbsp;1).\u003c/p\u003e\n\u003cp\u003eSpatio-temporal progression of the tremor hypocenters on 16 September 2005 (Fig.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e3\u003c/span\u003e) clearly indicates that the north and south clusters behave independently. Similar behavior is observed for the ETS episodes in 2003 and 2004 (Supplementary Fig.\u0026nbsp;5).\u003c/p\u003e\n\u003cp\u003eWe estimated the scalar moment and moment magnitude from the energy for each tremor detection. As in [\u003cspan class=\"CitationRef\"\u003e16\u003c/span\u003e], we observe a narrow normal distribution of magnitudes Mw\u0026thinsp;=\u0026thinsp;1.60\u0026thinsp;\u0026plusmn;\u0026thinsp;0.1, implying a log-normal distribution of scalar moments (Supplementary Fig.\u0026nbsp;6). Details on the moment and magnitude calculation are provided in Methods. Total moment released in each of the four clusters is given in Supplementary Table\u0026nbsp;1.\u003c/p\u003e\n\u003c/div\u003e"},{"header":"Discussion","content":"\u003cp\u003eAccurate depth determination of the individual LFEs that constitute tremor beneath Vancouver Island provide two constraints on their source: 1) their location relative to known subsurface structures and, 2) the width of the shear-zone that produces them.\u003c/p\u003e\n\u003cdiv id=\"Sec7\" class=\"Section2\"\u003e\n \u003ch2\u003eComparison of tremor locations with other geophysical observations\u003c/h2\u003e\n \u003cp\u003eConstraints on subsurface structure in the study region are provided by Lithoprobe seismic reflection profiling [\u003cspan class=\"CitationRef\"\u003e40\u003c/span\u003e], regional double-difference tomography incorporating LFE templates [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e], and receiver function studies [\u003cspan class=\"CitationRef\"\u003e42\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e43\u003c/span\u003e]. Figure \u003cspan class=\"InternalRef\"\u003e4\u003c/span\u003e shows tremor hypocenters from this study overlain on a northward extrapolation of the Lithoprobe 84\u0026thinsp;\u0026minus;\u0026thinsp;02 seismic reflection profile. Panel A includes Vp while panel B includes Poisson\u0026rsquo;s ratio. A dipping zone of quasi-parallel reflectors, dubbed the \u0026ldquo;E-layer\u0026rdquo;,[\u003cspan class=\"CitationRef\"\u003e40\u003c/span\u003e] was interpreted to represent underplated imbricate oceanic sediments and volcanics. However, tomographic studies constrain the P-wave velocities at the depth of the E-layer below southern Vancouver Island to be \u0026ge;\u0026thinsp;7 km/s (Fig. \u003cspan class=\"InternalRef\"\u003e4\u003c/span\u003e; [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e]) suggesting, instead, that this material is predominantly mafic [\u003cspan class=\"CitationRef\"\u003e44\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e45\u003c/span\u003e]. All of the E-layer, tremor and LFE template locations lie within a zone of unusually high Poisson\u0026rsquo;s ratio anomaly (~\u0026thinsp;0.28) that dips landward parallel to the slab, and is interpreted to indicate near-lithostatic fluid pressures in the tremor source region [\u003cspan class=\"CitationRef\"\u003e42\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e44\u003c/span\u003e].\u003c/p\u003e\n \u003cp\u003eOur tremor locations indicate quasi-planar, segmented layers in the plate boundary region just below the E-layer, at a depth near 39 km and are consistent with those of nearby LFE template locations [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e44\u003c/span\u003e] projected into the profile. Our observation that the E-layer lies above the active locus of seismic deformation (Fig.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e4\u003c/span\u003e) is contrary to reports from previous studies in this region [\u003cspan class=\"CitationRef\"\u003e13\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e25\u003c/span\u003e]. Although we cannot rule out the possibility that some tremor mapped away from the four principal patches represents true detections, almost all detections are relatively tightly constrained to layers less than 400 m thick.\u003c/p\u003e\n\u003c/div\u003e\n\u003cdiv id=\"Sec8\" class=\"Section2\"\u003e\n \u003ch2\u003eMafic underplating model for tremor\u003c/h2\u003e\n \u003cp\u003eBased on our locations, we interpret tremor beneath Vancouver Island to represent mafic underplating, wherein each LFE represents shear failure within mixed brittle-ductile deformation occurring in the top few hundreds of meters of crystalline oceanic crust (commonly referred to as layer 2A [\u003cspan class=\"CitationRef\"\u003e46\u003c/span\u003e]) of the subducting Juan de Fuca plate. This layer is expected to be rich in free fluids (to ~\u0026thinsp;4%; [\u003cspan class=\"CitationRef\"\u003e46\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e47\u003c/span\u003e]) as it undergoes active prograde metamorphism at tremor depths (e.g., [\u003cspan class=\"CitationRef\"\u003e47\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e48\u003c/span\u003e]). These fluids are produced by metamorphic reactions at lithostatic pore pressures (e.g., [\u003cspan class=\"CitationRef\"\u003e49\u003c/span\u003e]) and promote material weakening and localized deformation. We expect that fluid distribution within the layer 2A will be heterogeneous as governed by the distribution and persistence of fault-controlled fluid pathways connecting to the seafloor prior to subduction [\u003cspan class=\"CitationRef\"\u003e50\u003c/span\u003e], thus influencing where deformation is partly brittle versus where it is purely ductile. Simple shear induced through ongoing subduction causes the basaltic material to be continuously eroded leading to comminuted wear products with an increasingly anisotropic fabric, that are gradually plated onto the overriding lithosphere. Weaker volumes with higher water contents trapped by a highly anisotropic permeability are elongated by shear to produce the seismic reflectors that characterize the E-layer [\u003cspan class=\"CitationRef\"\u003e51\u003c/span\u003e], which is mostly aseismic (e.g., Fig.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e4\u003c/span\u003e). Localized areas where material transfer is occurring within the subduction zone may manifest the distinct tremor patches as seen in Figs.\u0026nbsp;\u003cspan class=\"InternalRef\"\u003e2\u003c/span\u003e and \u003cspan class=\"InternalRef\"\u003e3\u003c/span\u003e.\u003c/p\u003e\n \u003cp\u003eSupport for this interpretation is multifold. Based on analyses of multiple exhumed accretionary complexes,[\u003cspan class=\"CitationRef\"\u003e52\u003c/span\u003e] argued that exhumed, underplated basalt occurs as thin (\u0026le;\u0026thinsp;300 m) layers derived exclusively from layer 2A, consistent with the model interpretation laid out above. In a study of the Arosa zone, plate boundary rocks exhumed from tremor depths,[\u003cspan class=\"CitationRef\"\u003e53\u003c/span\u003e] noted that plate boundary slip occurs as frictional deformation in chlorite and talc schists surrounding blocks of metabasalt, consistent with our inference that tremor is hosted within layer 2A. Moreover, the remarkably coherent, coast-parallel distribution of tremor epicenters along the entire Cascadia margin seems unlikely were tremor associated exclusively with sediments given the variable sediment input along the margin. This observation provides additional support for a basaltic layer 2A origin of tremor beneath southern Vancouver Island. As previously noted by [\u003cspan class=\"CitationRef\"\u003e30\u003c/span\u003e], this epicentral distribution also generally mirrors the high coastal topography associated with warm subduction zone settings, and is readily explained if tremor is a manifestation of crustal underplating [\u003cspan class=\"CitationRef\"\u003e48\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e54\u003c/span\u003e]. Moreover, in the limited locations where crustal seismic profiling has been undertaken in the Cascadia forearc, notably from the Strait of Juan de Fuca through to central Vancouver Island [\u003cspan class=\"CitationRef\"\u003e40\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e55\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e56\u003c/span\u003e] and in central Oregon [\u003cspan class=\"CitationRef\"\u003e57\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e58\u003c/span\u003e], a highly reflective E-layer above the inferred slab has been identified. This suggests that it, like tremor, is present along the full Cascadia margin and is a key element of the subduction complex.\u003c/p\u003e\n \u003cp\u003eAlthough [\u003cspan class=\"CitationRef\"\u003e52\u003c/span\u003e] suggested a \u0026ldquo;peeling\u0026rdquo; of layer 2A in the transfer of metabasalt from subducting to overriding plates, the nature of tremor suggests an origin involving significant cataclasis [\u003cspan class=\"CitationRef\"\u003e59\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e60\u003c/span\u003e]. For example, [\u003cspan class=\"CitationRef\"\u003e16\u003c/span\u003e] observed log-normal distributions of LFE moments that that they argue can be explained via a model wherein shear failure at contacts between rigid grains jammed within a viscous channel generates tremor. In the current context, we interpret the granular and viscous elements of layer 2A to be associated with less altered tracts of metabasalt surrounded by a more intensely hydrated and overpressured matrix, respectively. The lognormal distribution of moments originates from the lognormal distribution of contact areas within jams expected as larger competent clasts are gradually broken down into smaller ones. As the clasts decrease in size, they become less prone to jamming, and we suggest that a scale-dependence set by layer 2A thickness contributes to the band-limitation (~\u0026thinsp;1\u0026ndash;10 Hz) of tremor [\u003cspan class=\"CitationRef\"\u003e16\u003c/span\u003e] and the limited range of magnitudes observed in [\u003cspan class=\"CitationRef\"\u003e16\u003c/span\u003e] and in this paper.\u003c/p\u003e\n \u003cp\u003eAs comminution proceeds, we expect increasing shear strain, ductile deformation, and gradual material transfer/transformation to the E-layer because of decreased density and strength imparted by the release of fluids [\u003cspan class=\"CitationRef\"\u003e61\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e54\u003c/span\u003e]. We note that tremor and LFE template hypocenters lie on average 2\u0026ndash;3 km below the base of the reflectivity that defines the E-layer in Fig.\u0026nbsp;5. We discount the possibility of location bias since the same velocity model[\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e] is used to locate hypocenters and migrate reflections. Rather, we argue that, at some point in the comminution and shearing process, a permeability anisotropy ``percolation\u0026rdquo; threshold is reached wherein fluids become segregated within horizontally contiguous \u0026ldquo;lenses\u0026rdquo; producing the pronounced reflectivity horizons and abrupt base that define the E-layer, perhaps through changes in dihedral angle as suggested by [\u003cspan class=\"CitationRef\"\u003e46\u003c/span\u003e]. Although our estimates of slip within tremorgenic volumes based on Kostrov strain significantly exceed those previously reported for tremor within the ETS zone more generally [\u003cspan class=\"CitationRef\"\u003e62\u003c/span\u003e], they nonetheless fall far short of the plate motion budget (~\u0026thinsp;3mm versus ~\u0026thinsp;4-5cm per ETS episode) indicating that ductile deformation must still dominate. It is likely then that the slow slip of ETS represents ductile shear persisting well into the lower reaches of the E-layer at steadily diminishing levels, both where tremor is well expressed and where it is not [\u003cspan class=\"CitationRef\"\u003e63\u003c/span\u003e].\u003c/p\u003e\n\u003c/div\u003e\n\u003cdiv id=\"Sec9\" class=\"Section2\"\u003e\n \u003ch2\u003eDistance-thickness calculation of the E-layer\u003c/h2\u003e\n \u003cp\u003eWe assess the feasibility of E-layer assembly over realistic time periods using calculations modified from [\u003cspan class=\"CitationRef\"\u003e40\u003c/span\u003e]. The E layer is roughly 100 km long and about 5 km thick beneath Vancouver Island, corresponding to a volume of 500 km\u003csup\u003e3\u003c/sup\u003e/km along strike. Roughly 1800 km of plate have been subducted beneath Vancouver Island over the last 40 Ma [\u003cspan class=\"CitationRef\"\u003e40\u003c/span\u003e]. If a thickness H of the basalt layer was eroded away during this time to form the E-layer, the eroded volume is 1800*H km\u003csup\u003e3\u003c/sup\u003e/km along strike. If the E-layer was formed through basal accretion of layer 2A, then 1800H\u0026thinsp;=\u0026thinsp;500, or H\u0026thinsp;=\u0026thinsp;0.28 km, in rough agreement with the thickness of the layer 2A pillow basalts [\u003cspan class=\"CitationRef\"\u003e64\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e65\u003c/span\u003e]. Moreover, the total relative displacement of the Juan de Fuca plate D and the thickness of the E-layer are consistent with the observed relation between displacement and thickness of crustal fault zones (Supplementary Fig.\u0026nbsp;7, [\u003cspan class=\"CitationRef\"\u003e66\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e67\u003c/span\u003e]).\u003c/p\u003e\n\u003c/div\u003e\n\u003cdiv id=\"Sec10\" class=\"Section2\"\u003e\n \u003ch2\u003eTremor as diagnostic of material transfer\u003c/h2\u003e\n \u003cp\u003eFinally, we note the additional occurrence of tectonic tremor beneath accretionary prisms in subduction zones such as Nankai [\u003cspan class=\"CitationRef\"\u003e68\u003c/span\u003e], where underplating is also implied, and at major strike-slip faults (such as the Alpine Fault [\u003cspan class=\"CitationRef\"\u003e69\u003c/span\u003e] or San Andreas Fault [\u003cspan class=\"CitationRef\"\u003e4\u003c/span\u003e]). We suggest that the occurrence of tremor in these environments, as in the deep plate boundary of subduction zones, may be diagnostic of granular flow and/or material transfer in zones of high pore-fluid pressure (e.g., [\u003cspan class=\"CitationRef\"\u003e70\u003c/span\u003e, \u003cspan class=\"CitationRef\"\u003e71\u003c/span\u003e]).\u003c/p\u003e\n\u003c/div\u003e"},{"header":"Methods","content":"\u003cdiv id=\"Sec12\" class=\"Section3\"\u003e\n\u003ch2\u003eDetection and Location\u003c/h2\u003e\n\u003cp\u003eWe employ the P- and S-waveforms for LFE templates 053, 065, and 070 at 5 stations to determine a) delays for S- and P-wave arrivals corresponding to the nominal template location, computed using the alignment procedures in [\u003cspan class=\"CitationRef\"\u003e72\u003c/span\u003e]; b) the splitting parameters that best reduce the S-wave particle motions on the two horizontal coordinates to rectilinear motion isolated to a single channel [\u003cspan class=\"CitationRef\"\u003e73\u003c/span\u003e]; and c) the expected P-polarity for P-waveforms on the vertical component. These quantities are used to normalize 24 hour-long waveforms for the 4 S-wave stations (KELB (2003)/KLNB (2004 and 2005), PGC, SILB, SSIB) and 1 P-wave station (SNB) employed in this study (stations KELB and KLNB differ in location by ~\u0026thinsp;40 m). PGC is designated as the time-stamp reference station (0 s delay) with template-dependent relative delays applied to the remaining 4 channels. Moreover, a single suite of station-specific splitting parameters is employed for all 3 templates to maintain the same S-waveforms (with different relative delays) for each case. Following Rubin and Armbruster, the normalized 24-hour waveforms are divided into 86396 4 s windows with 3 s overlap starting at midnight (PGC time) and cross-correlations are computed for all (4) pairs of S-stations with a common time stamp to maximum lags of +/- 0.4 s for a given template to mitigate against cycle skips (Supplementary Fig.\u0026nbsp;1). Thus, a tremor burst originating at the nominal template epicenter would register maximum correlation at 0 s lag for each pair, whereas non-zero lags characterize epicenters away from the template epicenter. The chosen +/- 0.4 s lag allows some overlap across the 3 template epicentral regions. A prospective detection is declared if 2 conditions are met ( [\u003cspan class=\"CitationRef\"\u003e8\u003c/span\u003e]; see Supplementary Note 1) relating to thresholds on the values of a) the 4 possible 3-station delay time circuits (i.e. | \u003cem\u003et\u003c/em\u003e\u003csub\u003e\u003cem\u003eij\u003c/em\u003e\u003c/sub\u003e+\u003cem\u003et\u003c/em\u003e\u003csub\u003e\u003cem\u003ejk\u003c/em\u003e\u003c/sub\u003e-\u003cem\u003eti\u003c/em\u003e\u003csub\u003e\u003cem\u003ek\u003c/em\u003e\u003c/sub\u003e| ), and the mean cross correlation coefficient. In the event of a prospective detection, the first principal component waveform of the aligned S-waveforms is cross-correlated with the single-station P-waveform and a detection is declared upon meeting a second correlation coefficient threshold. The lag at maximum correlation enables computation of an S-P time (and therefore hypocentral depth) and lags are again restricted to lie between +/- 0.4 s.\u003c/p\u003e\n\u003cp\u003eThe computations above are performed for two different frequency bands: a narrow band of 1.5-6 Hz like that employed by [\u003cspan class=\"CitationRef\"\u003e8\u003c/span\u003e], and a broader band of 1\u0026ndash;8 Hz which yields fewer detections, but reduced likelihood of cycle skips due to the wider range of frequencies represented. There is therefore the possibility for up to 6 duplicate detections per time stamp (2 frequency bands for each of 3 templates). We employ several thresholds and statistics (described in detail in the Supplementary Note 1) to cull the tremor catalogue to a maximum of 1 detection per time stamp to emphasize tremor hypocentral patterns but minimize scatter arising from false detections. We do not attempt to eliminate repeated arrivals across consecutive overlapping waveforms, thereby allowing for a continuously evolving tremor wavefield. Each detection is thus characterized by 4 S times and 1 P time allowing for location with only 1 degree of redundancy (5 constraints on 4 hypocentral parameters). This approach yields high relative precision in location (dependent upon the precision of the relative delays), but an absolute accuracy that is dependent upon the validity of the velocity model given that stations are dominantly to one side of the source region. The choice of a suitable P-wave station is a compromise in SNR between epicentral distance and relative P-to-S radiation (assessable from template waveforms) and is best met for this region by station SNB [\u003cspan class=\"CitationRef\"\u003e26\u003c/span\u003e].\u003c/p\u003e\n\u003cp\u003eWe obtain initial locations using Hypoinverse [\u003cspan class=\"CitationRef\"\u003e74\u003c/span\u003e] with a 1D velocity model based on the [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e] 3D model at this location. After culling detections (described in Supplementary Note 1), we determine final hypocenters using double-difference relocation in tomoDD [\u003cspan class=\"CitationRef\"\u003e75\u003c/span\u003e] with the 3D velocity model of [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e]. We also compare the resulting relocations using SSIB as the P-wave detection station instead of SNB. The relocations preserve a very similar relative pattern, although due to inaccuracy of the velocity model, absolute locations are shifted slightly deeper by about 1.5 km. However, relocation of the same template waveforms as [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e] using the method presented here shows good agreement with the [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e] study. Furthermore, our final tremor catalog locations show good agreement with the [\u003cspan class=\"CitationRef\"\u003e41\u003c/span\u003e] LFE family template locations, indicating that our method yields robust absolute locations (Supplementary Fig.\u0026nbsp;8). The final catalog represents 4,851 locations determined for a total of 13 days spread over 3 ETS episodes, as determined from the LFE catalogs of [\u003cspan class=\"CitationRef\"\u003e32\u003c/span\u003e].\u003c/p\u003e\n\u003c/div\u003e\n\u003cdiv id=\"Sec13\" class=\"Section2\"\u003e\n\u003ch2\u003eFalse Detections and Location Uncertainty\u003c/h2\u003e\n\u003cp\u003eAll detections scattered away from the main patches were visually inspected for cycle-skipping in the waveform alignment or false detections. Only about 15% of the scattered detections were deemed true detections, and most of these events occurred shallower than 37 km depth. We do not observe any significant tremor patches distributed in depth or anywhere significantly above or below about 39 km depth. We therefore cull the dataset further to include only the events between 37 and 39.5 km depth. Visual inspection of many of the waveforms contributing the to the detections composing the primary layer structures indicates that most of these detections do not suffer from cycle-skipping.\u003c/p\u003e\n\u003cp\u003eWe estimate the location uncertainty by comparing the locations of common timestamp detections between the narrow \u0026amp; broad band data in 3 different templates (each gives a different location measurement, resulting in 6 different locations to compare). Epicentral uncertainty is estimated to be about 250 m, and depth uncertainty about 300 m (Supplementary Figs.\u0026nbsp;3, 9). We note that the nominal horizontal and depth errors from Hypoinverse are 1.15 km and 2.3 km, respectively.\u003c/p\u003e\n\u003c/div\u003e\n\u003cdiv id=\"Sec14\" class=\"Section2\"\u003e\n\u003ch2\u003eMoment and Moment Magnitude Calculations\u003c/h2\u003e\n\u003cp\u003eFollowing [\u003cspan class=\"CitationRef\"\u003e8\u003c/span\u003e], we assign a consistent, coherent radiated energy metric for each 4 s, narrow-band S-detection window as\u003c/p\u003e\n\u003cdiv id=\"Equa\" class=\"Equation\"\u003e\n\u003cdiv id=\"FileID_Equa\" class=\"mathdisplay\"\u003e$$E\\left(t\\right)= \\frac{{S}_{A}\\left(t\\right){S}_{B}\\left({t}_{B}^{{\\prime }}\\right)+{S}_{A}\\left(t\\right){S}_{C}\\left({t}_{C}^{{\\prime }}\\right)+{S}_{B}\\left({t}_{B}^{{\\prime }}\\right){S}_{C}\\left({t}_{C}{\\prime }\\right)}{3}$$\u003c/div\u003e\n\u003c/div\u003e\n\u003cp\u003e,\u003c/p\u003e\n\u003cp\u003ewhere \u003cem\u003eS\u003c/em\u003e(\u003cem\u003et\u003c/em\u003e) is the S-wave seismogram, \u003cem\u003eA\u003c/em\u003e, \u003cem\u003eB\u003c/em\u003e, and \u003cem\u003eC\u003c/em\u003e denote stations PGC, SILB, SSIB, and \u003cem\u003et\u003c/em\u003e\u0026rsquo; is the time offset between the tremor arrival at stations \u003cem\u003eB\u003c/em\u003e or \u003cem\u003eC\u003c/em\u003e and station \u003cem\u003eA\u003c/em\u003e. To assign magnitudes from energy, we identify those 4 s windows (a total of 391) that correspond to matched filtered LFE detections with magnitudes from the catalog of [\u003cspan class=\"CitationRef\"\u003e32\u003c/span\u003e], and perform a power-law regression to establish the appropriate calibration for moment magnitude:\u003c/p\u003e\n\u003cdiv id=\"Equb\" class=\"Equation\"\u003e\n\u003cdiv id=\"FileID_Equb\" class=\"mathdisplay\"\u003e$${M}_{w}=0.238\\text{*}{log}_{10}\\left(E\\right)+0.319,$$\u003c/div\u003e\n\u003c/div\u003e\n\u003cp\u003eyielding the scalar moments for all available S-wave windows:\u003c/p\u003e\n\u003cdiv id=\"Equc\" class=\"Equation\"\u003e\n\u003cdiv id=\"FileID_Equc\" class=\"mathdisplay\"\u003e$${M}_{0}={10}^{(1.5{M}_{w} + 10.7)}$$\u003c/div\u003e\n\u003c/div\u003e\n\u003cp\u003e.\u003c/p\u003e\n\u003c/div\u003e"},{"header":"References","content":"\u003col\u003e\n\u003cli\u003eG. C. Rogers and H. 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Unsworth, \u0026quot;Cascadia low frequency earthquakes at the base of an overpressured subduction shear zone,\u0026quot; \u003cem\u003eNature Communications, \u003c/em\u003evol. 11, no. 3874, 2020. \u003c/li\u003e\n\u003c/ol\u003e"}],"fulltextSource":"","fullText":"","funders":[],"hasAdminPriorityOnWorkflow":false,"hasManuscriptDocX":true,"hasOptedInToPreprint":true,"hasPassedJournalQc":"","hasAnyPriority":true,"hideJournal":true,"highlight":"","institution":"","isAcceptedByJournal":false,"isAuthorSuppliedPdf":false,"isDeskRejected":"","isHiddenFromSearch":false,"isInQc":false,"isInWorkflow":false,"isPdf":false,"isPdfUpToDate":true,"isWithdrawnOrRetracted":false,"journal":{"display":true,"email":"[email protected]","identity":"researchsquare","isNatureJournal":false,"hasQc":true,"allowDirectSubmit":true,"externalIdentity":"","sideBox":"","snPcode":"","submissionUrl":"/submission","title":"Research Square","twitterHandle":"researchsquare","acdcEnabled":true,"dfaEnabled":false,"editorialSystem":"","reportingPortfolio":"","inReviewEnabled":false,"inReviewRevisionsEnabled":true},"keywords":"","lastPublishedDoi":"10.21203/rs.3.rs-3909443/v1","lastPublishedDoiUrl":"https://doi.org/10.21203/rs.3.rs-3909443/v1","license":{"name":"CC BY 4.0","url":"https://creativecommons.org/licenses/by/4.0/"},"manuscriptAbstract":"\u003cp\u003eTremor is a low-amplitude seismic signal that usually temporally coincides with episodic slow fault slip at plate boundaries worldwide. Since the discovery of tremor in Cascadia, significant effort has been devoted to understanding its relationship to slow slip. However, its source mechanism has been widely debated, owing in large part to the challenge of locating sources accurately in depth. We assemble a tremor catalog of 4,851 events for a\u0026thinsp;~\u0026thinsp;10 X 20 km\u003csup\u003e2\u003c/sup\u003e area on southern Vancouver Island from slow slip episodes in 2003\u0026ndash;2005 using a cross-station detection method adapted from previous studies, which we extend to use both P- and S- waves, thereby recovering accurate depths. Tremor occurs in distinct, quasi-planar clusters in the plate boundary region at a depth near 39 km, just beneath a layer of high reflectivity and within a zone of elevated Poisson\u0026rsquo;s ratio. We interpret this tremor to represent mafic underplating, wherein shearing generates tremor and continuously erodes basaltic material of the upper few hundred meters of the oceanic crust. Comminuted basalt with an increasingly anisotropic fabric is gradually plated onto the overriding lithosphere to form the highly reflective layer. Localized areas of material transfer within the subduction zone may manifest the distinct tremor clusters.\u003c/p\u003e","manuscriptTitle":"Tectonic tremor: the chatter of mafic underplating beneath southern Vancouver Island?","msid":"","msnumber":"","nonDraftVersions":[{"code":1,"date":"2024-02-09 18:52:49","doi":"10.21203/rs.3.rs-3909443/v1","editorialEvents":[{"type":"communityComments","content":0}],"status":"published","journal":{"display":true,"email":"[email protected]","identity":"researchsquare","isNatureJournal":false,"hasQc":true,"allowDirectSubmit":true,"externalIdentity":"","sideBox":"","snPcode":"","submissionUrl":"/submission","title":"Research Square","twitterHandle":"researchsquare","acdcEnabled":true,"dfaEnabled":false,"editorialSystem":"","reportingPortfolio":"","inReviewEnabled":false,"inReviewRevisionsEnabled":true}}],"origin":"","ownerIdentity":"a34f578c-a5b5-4aac-932b-a3e1c08ed985","owner":[],"postedDate":"February 9th, 2024","published":true,"recentEditorialEvents":[],"rejectedJournal":[],"revision":"","amendment":"","status":"posted","subjectAreas":[{"id":28460004,"name":"Earth and environmental sciences/Solid Earth sciences/Seismology"},{"id":28460005,"name":"Earth and environmental sciences/Solid Earth sciences/Tectonics"}],"tags":[],"updatedAt":"2025-02-25T16:11:59+00:00","versionOfRecord":[],"versionCreatedAt":"2024-02-09 18:52:49","video":"","vorDoi":"","vorDoiUrl":"","workflowStages":[]},"version":"v1","identity":"rs-3909443","journalConfig":"researchsquare"},"__N_SSP":true},"page":"/article/[identity]/[[...version]]","query":{"redirect":"/article/rs-3909443","identity":"rs-3909443","version":["v1"]},"buildId":"qtupq5eGEP_6zYnWcrvyt","isFallback":false,"isExperimentalCompile":false,"dynamicIds":[84888],"gssp":true,"scriptLoader":[]}