In this section, we identify the data sources that were integrated for analyses of recorded historical earthquakes in different tectonic settings, including the Cascadia convergent zone. The compilations of “seismic” events in the Cascadia convergent zone also include catalogs of recent Episodic Tremor and Slip (ETS) events. Transects for previously reported upper-plate strain data in the Cascadia convergent zone are shown and described. Comparisons between the modern strain data, the historical seismic record, recent ETS events from the Cascadia convergent zone are presented in the Results section of this article. In general, the worldwide earthquake catalog data [36] [37] [38] were imported into a relational database, and the various catalogs were joined by finding information in common between the datasets, mostly the catalog number. In some instances, the epoch, location, and magnitude were used to relate datasets. Datasets, such as the displacement data reported by Wells and Coppersmith [39] and others [40] [41] [42] [43] could then be joined to known events and Moment Tensor event information. Datasets were joined to the various catalog data by adding a USGS catalog number to the datasets. This method joined the offset data to the catalog event, which was tied to Flinn-Engdahl regions [44] , and from there to tectonic boundary type. The USGS event was also joined to the Centroid Moment Tensor catalog and the Radiated energy catalogs. Earthquakes in the catalog could be characterized using the Flinn-Engdahl region [44]. Each Flinn-Engdahl region was then ascribed to the dominant tectonic setting of the region. The mapped tectonic boundaries were then grouped under simplified tectonic boundaries, as shown in Table 1. The completed database allowed for the integration of various seismological and geological observations. Database script was written to produce the tables and figures used in this article.

Earthquake catalogs can be difficult to use for tectonic analysis. For example, we focus on magnitude five and larger earthquakes in some analyses, since these are the most reliably detected by worldwide networks. Local networks, the qualities of which vary regionally, could put biases in the data, so we favor the magnitude five and larger events. We also focus on shallow- and intermediate depth events (0 - 60 km depth; [45] , p. 30). Deeper events (>60 km) might not be

clearly related to the surface tectonic setting, but they do have an application to mapping the descending slab or lower-plate. When using the Pacific Northwest Seismic Network events [46] , volcanic earthquakes were excluded from the compilations used in this article, following the determination made by Pacific Northwest Seismic Network (which covers the central and southern Cascadia margin). If the earthquake is in their “volcanic” source list, we ignore that event. It is not 100% effective, but in bulk, it seems to eliminate the hotspots associated with the Cascade volcanoes. Lastly, we do not consider Richter-like magnitudes to be the best measure of earthquake energy. Where possible, we use the scalar moment from the Harvard CMT, or the IRIS radiated energy. We prefer radiated energy since it makes fewer assumptions

In this paper, we compile and integrate the following datasets:

· Convergent Margin Strain studies [22] [24] [26]

· PNSN Earthquake locations [46]

· PNSN First-Motion Studies [46]

· PNSN Tremor events [46]

· SPUD [48] and CHOY [49] catalogs of Radiated Seismic Energy

GPS baseline strains, i.e., baseline shortening or lengthening, between 1200 paired adjacent GPS stations are well documented in the Cascadia margin [24]. Details on the methods of strain analyses are given elsewhere [22]. For the purposes of convergent strain analyses in this paper, we use selected margin perpendicular (west-east) GPS station transects, which are normal to the major linear structural elements or deformation bands in the Cascadia convergent zone (Figure 1 and Figure 2). The density of GPS base stations and previously published baseline strain transects are shown in Figure 3. Baseline strain, averaged over decadal intervals, is normalized to annual rates of shortening (negative strain) or lengthening (positive strain) and plotted between GPS station endpoints.

The University of Washington earthquake catalog [46] was used for compilations of a total of about 106,000 earthquakes (−2.5 ≤ M ≤ 6.8) in the Cascadia convergent zone. Earthquake data compilations include event dates, hypocenter locations, magnitudes, and first-motions. The maps in this article omit earthquake events identified in the UW catalog as belonging to volcanic activity associated

with the various Cascade volcanoes. This does not eliminate all the probable volcanic earthquakes. Hypocenter data is shown in two ways: One is a traditional circle at the location of the earthquake; the other is as earthquake density. For earthquake density, the number of earthquakes within a small area is represented by a contour map.

For Cascadia episodic tremor and slip (ETS) events the University of Washington catalog [46] was used. Over 400,000 ETS events from the Cascadia subduction zone were compiled for this article. As with earthquake epicenters, the ETS events are contoured by density to show areas of relative concentration. The depths of ETS events are not well constrained [29] [50] , but for the purposes of this article, they are assumed to lie near or within the inter-plate interface. Reported features of some ETS events include periodic displacements at annual or sub-annual time scales, including for example, an apparent westward movement of station ALBH in Victoria, British Columbia, Canada [51]. This apparent movement is relative to station DRAO in the Canadian Rockies and assumes DRAO is fixed and unmoving. The NASA/JPL data using an Earth-Centered, Earth-Fixed (ECEF) solution suggests both stations ALBH and DRAO are moving toward the southwest. DRAO (16.6 mm/yr) is moving faster than ALBH (11.3 mm/yr). The apparent relative periodic displacements between the two stations are still resolved, but they are not explained. In this article, we do not address ETS periodic movements but focus on discrete events. Details on processing and selecting ETS positions are given by Wech [29] [50].

Surface GPS stations can be used to look at the variation in the vertical component of GPS station motion [24]. Though not recording vertical strain, the GPS station velocity data do demonstrate vertical deformation at the regional scale. The vertical velocity data were obtained from the Plate Boundary Observation stations, as reported by UNAVCO [52]. Preliminary test and calibrations of the regional vertical velocity data were made using previously published vertical displacement trends from 50-year geodetic releveling surveys [53].

In the preceding sections, we have shown that convergent margins are seismologically different from other margins. In the remainder of this article we compile reported modern strain and recent seismicity data for the Cascadia margin region. These dataset compilations and analyses are used to show how the primary seismogenic structure in the Cascadia convergent zone is expressed and how it could control great earthquake energy or magnitude during major mega-thrust ruptures. The Cascadia margin differs from some other subduction zones on the basis of 1) locally oblique plate convergence and 2) low angles of oceanic plate dip (5˚ - 12˚) as shown respectively in Figure 1 and Figure 2. However, the central portion of the Cascadia margin (Figure 1) is relatively uncomplicated (straight). It has a well-dated paleo-record of coseismic coastal subsidence and corresponding nearfield tsunami excitation from major mega-thrust ruptures [60] [61]. Of particular importance to the analyses presented here, the Cascadia margin is well-instrumented for GPS station baselines [22] [24] and broadband seismograph stations [46]. However, the Cascadia margin has not experienced a great earthquake in historical time (pre-European contact), thus motivating studies of predicting potential seismic energy distributions from a future major mega-thrust rupture.

Measured convergent strains measured in GPS baselines in the Cascadia margin range from 10⁻⁷ a⁻¹ to 10⁻⁹ a⁻¹ [24]. Representative east-west or margin parallel strain transects are shown in Figure 5. Modern crustal shortening dominates across the Coast Ranges in Vancouver Island, British Columbia, the Olympic Range in northwest Washington, the Northern Coast Range in southwest Washington and Oregon, and the Siskiyou/Klamath Coast Ranges in southwest Oregon and northwest California. Convergent annual strain rates in the Coast Ranges (generally −10⁻⁸ a⁻¹ to −10⁻⁷ a⁻¹) represent modern crustal shortening due to underlying inter-plate coupling and resulting shear stress transfer through the upper-plate to manifest as strain at the upper-plate surface. As will be addressed below, infrequent ETS events at or near the inter-plate interface suggest episodic slip between the plates under the Coast Ranges. Furthermore, changing strain rates in some transects [24] suggest transient strain propagation across the coupled zone, possibly related to components of aseismic slip across the coupled zone.

A landward band of little to no convergent strain (Figure 5) approximately coincides with the large forearc valleys (Puget Trough and Willamette River Valley) which lie between the Coast Ranges and the Cascade Range in the central Cascadia margin (Figure 1 and Figure 2). This narrow band (20 - 50 km in width) is interpreted to represent a very-weakly or intermittently decoupled interface between the underlying Jan de Fuca plate and the overriding North American plate [22] [31]. As well be shown below, the maximum density of ETS events are localized in this narrow band of little to no modern convergent strain. A more landward band of anomalous high convergent strain (annual strain rates −10⁻⁸ a⁻¹ to −10⁻⁷ a⁻¹) is locally present along the western side of the Cascade volcanic arc in Washington and northern Oregon and in at least one transect in southern Oregon (Figure 5). Crustal shortening in this narrow band suggests inter-plate coupling under the volcanic arc where the North American plate is slightly thickened [30]. Where the continental plate thins, landward (east) of the present volcanic arc the east-west GPS baselines are generally characterized by no measurable change in strain (<10⁻⁹ a⁻¹) or by local extension. These baselines have been interpreted to represent decoupling between the upper- and lower-plates, though one transect (Long Beach transect; Figure 10) in central Washington demonstrated significant crustal shortening on the east side of the Cascade Range. As will be addressed below, this anomalous zone of convergent strain on the east side of the Cascade Range in Washington coincides with locally historical seismicity in the upper plate.

Plots of hypocenter location and magnitude for deeper earthquakes (>30 km depth) in the Pacific Northwest region, including the Cascadia convergent zone, and shallow earthquakes (<30 km depth) in the upper-plate are presented in Figure 6. The cataloged earthquakes (106,000 in number) recorded between March 1969 and May 2019 demonstrate small-magnitude (M ≤ 6.8) releases of elastic strain energy in either the inter-plate interface or the descending slab (deeper events in Figure 6(a)), or the upper-crust (shallow events in Figure 6(b)) and offshore in the Gorda plate (Figure 1). The deeper inter-plate or descending slab events are concentrated in northwest Washington, where the Cascadia margin substantially bends, possibly introducing confining forces on the descending slab. Relatively little deep (inter-plate) seismicity is associated with the remainder of the central Cascadia margin (southwest Washington and Oregon), which host the most complete records of major mega-thrust rupture data, including: wide-spread coseismic coastal subsidence [59] [61] [62] , nearfield

tsunami inundation [63] , and coseismic paleoliquefaction [25] [64]. The historic inter-plate seismic record does not serve as an indicator of past major mega-thrust ruptures in the Cascadia convergent zone, but it does help define the extent of the primary seismogenic structure.

The distributions of cataloged smaller-magnitude earthquakes (M ≤ 6.8) compare well to modern convergent strain in the upper- and lower-plates (Figure 5 and Figure 6(b)). Relatively dense distributions of shallow earthquake events (≤30 km depth) indicate stress/strain resulting from convergence of the Juan de Fuca and North American plates and extending to about 400 - 500 km from the buried trench (also referred to as the Cascadia deformation front) in Washington and northern Oregon. The frequency and magnitude of historical seismicity decrease dramatically in southern Oregon, where infrequent upper-plate seismic events reach distances of 300 km from the buried trench. As will be shown below, approximate distributions of episodic tremor and slip (ETS) events compare favorably to margin-parallel bands of upper-plate shortening (moderately-coupled inter-plate zones) and upper-plate stability or extension (locally decoupled inter-plate zones).

In summary, the convergence-related shallow earthquakes (<30 km depth) in the upper plate of the Cascadia convergent zone are concentrated in western Washington, northwest Oregon, and southernmost British Columbia (Figure 6(b)). The upper-plate seismicity in Washington extends in relatively-high abundance across the Coast Ranges, Forearc valley, Cascade Range, and east of the Cascade Range to approximately 500 km landward of the buried trench. In contrast, upper-plate convergence-related seismicity in central and southern Oregon is sparse and only extends to about 300 km distance landward from the buried trench. The pattern of decreasing upper-plate seismicity from northern Oregon to southern Oregon is similar to decreasing upper-plate strain rates (crustal shortening) from northern Oregon to southern Oregon in the central Cascadia margin (Figure 5). It is not known whether the southward gradients of decreasing convergent strain and seismicity in the central Cascadia margin are due to 1) changes in plate convergence direction, 2) relative strengths of inter-plate coupling, and/or 3) upper-plate rigidity/strength characteristics (see further discussion below). The high abundance of shallow upper-plate earthquakes in the northern part of the central Cascadia margin differs from most other subduction zone margins (e.g., Alaska, Japan, Chile) where most of the earthquakes are in the inter-plate or subducting slab regions [37].

Two striking features of the Cascadia primary seismogenic structure in the central margin (Juan De Fuca Plate segment) area are 1) the relatively constant widths of the uplifted Coast Ranges (Figure 1) and the corresponding broad band of ETS events (Figure 1 and Figure 7(a)) and 2) the substantial differences in earthquake frequency between Washington/northern Oregon and central/southern Oregon (Figure 6 and Figure 7(b)). Some of the upper- and lower-plate seismicity in Northwest Washington could result from the sharp bend (nearly 45˚) of the plate margin in that area (Figure 1) and the associated confining forces associated with subduction and inter-plate coupling there. Whereas the Olympic Coast Range reaches greater elevations than the more southward Coast Ranges in Washington and Oregon [25] the approximate widths of the uplifted ranges remain relatively similar (~100 km) along the length of the central margin. The band of ETS events only slightly widens in northern Washington and southern Oregon, relative to central Oregon, suggesting similar widths of inter-plate coupling seaward (west) of the forearc valleys (Figure 7(a)). So what could give rise to the abundance of upper-plate and inter-plate earthquakes in the forearc valley, Cascade Range, and especially the retro-arc areas Washington and northernmost Oregon relative to the southern half of the central and southern Oregon (Figure 6 and Figure 7(b))? The answer could be indicated by the changing convergence strain rates in the Cascade Range: locally high-rates in the North Cascades (north of the Columbia River) and generally-low rates in the South Cascades (south of the Columbia River) (Figure 1 and Figure 5). The inter-plate recoupling is apparently greater under the North Cascade Range than the South Cascade Range, thereby yielding more convergence strain and seismicity in Washington and northernmost Oregon, relative to central and southern Oregon. Do the apparent differences in inter-plate recoupling signify greater upper-plate thickness/rigidity under the North Cascades than the South Cascades? Is it only a coincidence that the retro-arc Columbia River in Washington cuts across (is antecedent) through the volcanic arc at about the same position as the change in upper-plate seismicity and modern strain rates? Or are all three conditions related to changes in upper-plate thickness and inter-plate recoupling at about the latitude of Portland, Oregon?

Less-striking along-margin variations in the Cascadia primary seismogenic structure include the relative densities of ETS events at the landward (eastern) side of the ETS band (Figure 7(b)). Slightly higher-densities are shown in northern Washington and southern Oregon/northern California, relative to central Oregon. The higher-density ETS contours that are presented in this article are not short ellipses, as suggested by Bodmer, Toomey [66] , but rather are substantially elongated along-margin. Bodmer, Toomey [66] propose that the concentrated ETS clusters are associated with relatively greater inter-plate coupling, but modern convergence strain rates in the upper-plate (Figure 5) do not support that hypothesis. The greatest density of ETS events occurs in the southern Cascadia margin, the Gorda Plate segment, where upper-plate convergence stain rates are moderate by comparison to the central Cascadia margin [24]. Bodmer, Toomey [66] also suggest that the northern high-density clusters of ETS events could represent along-margin segmentation of the inter-plate coupled zone, thus limiting mega-thrust rupture lengths. The rupture boundary proposed by Bodmer, Toomey [66] is between the southern end of the Olympic Range and the Northern Coast Range in southwest Washington and Oregon (Figure 1). Ironically, the best paleo-seismic records of along-margin ruptures occur across that proposed boundary, with at least 5 out of the last 6 major mega-thrust ruptures, during the last ~2.6 ka, crossing from the central Washington coast to the northern Oregon coast [23] [61] [63] , and 3 out of the last 4 major mega-thrust ruptures, during the last 1.3 ka, crossing from the northern Washington coast to the central Oregon coast [67]. The apparent along-margin clustering of ETS events at the eastern margin of the ETS band does not apparently reflect either 1) the strength of inter-plate coupling or 2) major mega-thrust rupture lengths from the coupled zone located seaward (west) of the forearc valleys in the central Cascadia margin.

Comparing modern Cascadia upper-plate strains to those reported for the 2011 Tōhoku, Japan great earthquake (Mw~9) we would expect only about 100 years of present convergent strain rates in the central Cascadia margin to build up to Mw~9 energy-equivalent levels [24]. If we consider ~450 years to be the mean recurrence interval of major mega-thrust ruptures (Mw 9) and associated margin-long paleotsunami inundations in the Cascadia convergence zone [60] [68] , then less than 50 years could be sufficient to accumulate sufficient elastic strain to produce a lower-magnitude (M 8.0) event, at the observed strain rates. However, no great earthquakes (Mw ≥ 8.0) have occurred during the 300 years since the last major mega-thrust rupture at AD 1700 [68]. Some smaller-magnitude earthquakes (M ≤ 6) have occurred in or near the coupled-plate interface of the central Cascadia margin (Figure 5, Figures 9-11) and ETS events are mapped throughout most of the interpreted coupled zone (Figure 7(a)), though maximum frequency distributions are localized (Figure 7(b)). Such “seismic” events serve to release stress and accumulated strain in the convergent zone. Most recently, changes in observed strain rates have been reported for the central Cascadia margin [24] , suggesting aseismic propagation of strain across the western portion of the convergent zone. Presumably, some aseismic slip is occurring throughout the coupled plate interface, but enough shear stress is transmitted into the upper-plate to account for the modern strain accumulations monitored there. Furthermore, there has been relatively little reversal of vertical deformation associated with the mega-fold formed between the 1^st and 2^nd zero-isobases near the coast (Figure 2), as developed early in the present interseismic interval [24]. The complex relations noted above raise a very important issue. Such apparent historic and modern strain release processes, both seismic and aseismic, have not triggered a major mega-thrust rupture or great earthquake in the Cascadia margin, since the last great earthquake AD 1700. Such mega-thrust ruptures (hundreds of kilometers in length) are known to occur several century recurrence intervals throughout the Cascadia margin [60] [61] [62] [69] [70]. What then are the mechanisms that lead to a coseismic release of the accumulated elastic strain during the major mega-thrust ruptures? More generally, what triggers the major mega-thrust ruptures? The answer(s) to this question might be complex, involving both net accumulated strain and coincidences of multiple transient strain events. At the present time, the specific mechanisms that lead to major mega-thrust ruptures in the Cascadia primary seismogenic structure are not known.

We attribute the margin-parallel orientations of the large linear structural elements in the central Cascadia margin (Figure 1 and Figure 2) to across-margin gradients of upper-plate thickness and associated plate strength. Maximum upper-plate thickness under the Cascade volcanic arc [30] , sufficiently deflects the underlying Juan de Fuca oceanic plate to a depth of initial melting that feeds the arc volcanism, thereby further loading and thickening the upper-plate along the north-south volcanic arc. Slight thinning of the upper plate to the west (seaward) of the volcanic arc likely weakens inter-plate coupling along the forearc valleys, including the Puget Trough and Willamette Valley, both aligned north-south [31]. Further west (seaward) the shallowing dip angle (7˚ - 10˚) of the subducting Juan de Fuca plate strengthens the inter-plate coupling, resulting in low-angle thrusting and/or underplating that has uplifted the Coast Range, generally striking north-south [23] [25]. Continued thinning and associated weakening of the upper plate to the west (seaward) towards the continental shelf, has permitted thrust faults and associated folds to extend to the upper-plate surface. The terminal thinning of the upper-plate occurs at the trench where pull-down, tectonic erosion, and/or burial have obscured the seaward edge of the upper plate. However, the buried trench in the central Cascadia margin is generally aligned north-south. Therefore, the north-south oriented contours of upper-plate thickness, which are inherited structural controls, have guided the 1) lower-plate melting, 2) effective inter-plate coupling, and 3) upper-plate inelastic strain to yield the north-south striking (margin parallel) topographic bands between the buried trench and the volcanic arc in the central Cascadia margin.

The unusually shallow dip angles (5˚ - 12˚) of the subducting oceanic plate segments in the Cascadia margin (Figure 2) raise an interesting conundrum. Are the oceanic plate segments subducting under the over-riding North American continental plate or is the North American plate over-thrusting over the young and buoyant oceanic plate segments? Though likely important in terms of tectonic driving forces, we leave this question unanswered in this article, as the spatially variable inter-plate coupling and upper-plate strain accumulation in the Cascadia convergent zone are self-evident, regardless of the larger tectonic driving forces. As will be shown later in this article the underlying Juan De Fuca plate separates and descends below the North American plate to the east (landward) of the Cascade volcanic arc. That location of separation also demarks the landward-most possible extent of inter-plate coupling in the Cascadia convergent zone. Landward of the location of plate separation any localized convergent strain propagation further landward (east) would be transmitted entirely through the upper plate.

Substantial annual rates of convergent horizontal strain (10⁻⁸ a⁻¹ to 10⁻⁷ a⁻¹) are mapped across the full widths of the Coast Ranges to the forearc valleys in the Cascadia convergent zone (Figure 5). Equivalent annual strain rates are assumed to extend across the continental shelf, though a lack of GPS stations located seaward of the coastline precludes direct strain measurements across the shelf. Such high rates of annual convergent strain are also measured locally under the Cascade volcanic arc in Washington and northernmost Oregon, demonstrating effective inter-plate recoupling landward of the partially decoupled zone under the Puget and Willamette forearc valleys (Figure 2 and Figure 5). The inter-plate recoupling is interpreted to reach about 300 km distance from the buried trench in the southern Cascades of Oregon and 350 km distance from the trench in the northern Cascade terrain(s) in Washington and northernmost Oregon, as based on locally-high annual rates of convergent strain (10⁻⁸ a⁻¹). At greater landward distances the convergent strain rates diminish (10⁻⁹ a⁻¹) and are then replaced by locally high rates of west-east extension and/or north-south shortening [22] [24]. The across-margin transition from dominant west-east (convergent) shortening horizontal stain to dominant north-south shortening strain is used here to demark the landward extent of the primary seismogenic structure in the Cascadia convergent zone (Figure 13).

Horizontal strains, measured normal to plate margin orientation, represent modern convergent strain accumulation in the upper-plate of convergent margins. In this section such convergent margin strains were summarized for the central Cascadia margin, where the most complete strain data have been presented for the Cascadia subduction zone [24] and where major metropolitan centers occur in forearc valleys between the uplifting Coast Ranges and the volcanic arc (Figure 1, Figure 13). Modern annual convergent strains of 10⁻⁸ a⁻¹ to 10⁻⁷ a⁻¹ occur 1) regionally across the uplifting Coast Ranges to distances of 200 km from the buried trench and 2) then locally, with recoupling under the Cascades volcanic arc, to distances of up to 300 km and 450 km, respectively, in the southern and northern Cascades (Figure 5). Small, but significant, annual convergent strain rates of 10⁻⁹ a⁻¹ occur to landward distances of nearly 500 from the buried trench in Washington and northernmost Oregon, where the dominant strain changes from convergent (west-east) to margin parallel (north-south). For the purposes of simplification, two regional bands of modern horizontal convergent strain accumulation, including high rates and low rates, respectively are shown to extend to 300 km and 500 km in distances landward from the buried trench in the central Cascadia margin (Figure 13).

A complementary data set, to the modern horizontal strain, that demonstrates the broad extent of the Cascadia primary seismogenic structure (Figure 13) is the modern vertical deformation or GPS vertical velocities, as previously shown in Figure 8. A broadband (~100 km width) of general uplift occurs along the Olympic, North, and South Coast Ranges. A band of relative subsidence occurs landward of the Coast Ranges, but it varies substantially in across-margin width. The subsidence band is roughly centered along the Cascade volcanic arc, but it widens greatly from southern Oregon (~50 km in width) to northern Oregon (~150 km width) to Washington (~300 km width). It is not presently known whether the band of subsidence represents a continuation of across-margin mega-folding or includes a component of upper-plate pull-down by interaction with the descending plate in the inter-plate recoupled zone, located against the Cascade volcanic arc. The great width of the subsidence band in Washington is presently unexplained, but it does represent a broad region of modern elastic deformation, which coincides with the extended zone of convergent horizontal strain in Washington (Figure 5). Taken together, the along-margin bands of modern relative uplift and subsidence, and the across- margin distributions of convergent horizontal strain map out the potential extent (Figure 13) of potential seismic energy release during a major mega-thrust rupture in the Cascadia convergence zone.