Landslide mobility can vastly amplify the consequences of slope failure. As a compelling example, the 22 March 2014 landslide near Oso, Washington (USA), was particularly devastating, traveling across a 1-km+-wide river valley, killing 43 people, destroying dozens of homes, and temporarily closing a well-traveled highway. To resolve causes for the landslide’s behavior and mobility, we conducted detailed postevent field investigations and material testing. Geologic and structure mapping revealed a progression of geomorphological structures ranging from debris-flow lobes at the distal end through hummock fields, laterally continuous landslide blocks, back-rotated blocks, and finally colluvial slides and falls at the landslide headscarp. Primary structures, as well as stratigraphic and vegetation patterns, in the landslide deposit indicated rapid extensional motion of the approximately 9 × 10 6 m 3 source volume in a closely timed sequence of events. We identified hundreds of transient sand boils in the landslide runout zone, representing evidence of widespread elevated pore-water pressures with consequent shear-strength reduction at the base of the slide. During the event, underlying wet alluvium liquefied and allowed quasi-intact slide hummocks to extend and translate long distances across the flat valley. Most of the slide material itself did not liquefy. Using geotechnical testing and numerical modeling, we examined rapid undrained loading, shear and collapse of loose saturated alluvium, and strong ground shaking as potential liquefaction mechanisms. Our analyses show that some layers in the alluvium can liquefy when sheared, as could occur with rapid undrained loading. Simultaneous ground shaking could have contributed to pore-pressure generation as well. Two key elements, a large and rapid failure overriding wet liquefiable sediments, enabled the landslide’s high mobility. Basal liquefaction may enhance mobility of other landslides in similar settings.

We begin by providing a brief background on landsliding in the region and at the site of the Oso landslide. Following a description of our methods, we present our data and observations, and then we analyze these data to distinguish the major components of the landslide from one another. We subsequently use these components to interpret and describe a likely sequence of events, thus separating observations from interpretations. We then analyze our proposed mobility mechanisms involving liquefaction processes. Finally, we conclude with a discussion focused on the implications of the mobility of the Oso landslide for other similar settings, and a comparison of the various mechanisms put forth to explain mobility at Oso.

Here, we used postevent field observations combined with subsurface information and materials testing to decipher the geologic materials and processes of the 2014 Oso landslide, and to clarify how the sequence of events led to the landslide’s enhanced mobility. Our research was aided by fundamental observations obtained via a detailed field mapping effort performed over a 3-yr-period following the landslide. Our investigation focused on understanding the Oso landslide’s mobility, rather than on the causes for its initiation. However, we expect that our findings will provide critical information for subsequent studies that test initiation models.

The Oso landslide occurred at 10:36 a.m. local time (Pacific Daylight Time) on a sunny Saturday morning that followed both anomalously high March and 4 yr cumulative rainfall ( Henn et al., 2015 ; Iverson et al., 2015 ). The landslide initiated from a 180-m-tall terrace of Pleistocene glacial deposits located on the north side of the North Fork Stillaguamish River ( Dragovich et al., 2003 ). The source area was 760 m in length, but the landslide traveled over 1.5 times this distance (as measured from the toe to the most distal deposit), crossing the entire North Fork Stillaguamish River valley, and inundated 0.87 km 2 of the relatively flat valley floor including part of State Route 530 (SR530; Fig. 2 ). As a result, the Steelhead Haven neighborhood ( Fig. 2 ), located on the south side of the North Fork Stillaguamish River, was overrun by debris and destroyed. Remarkably, 11 persons in the neighborhood survived the rapid onslaught of mud and debris and were rescued from distal edges of the deposit ( Quistorf, 2014 ).

The episodes of previous landsliding, combined with ongoing forestry practices ( Everett and Casey, 2014 ), led to a variable forest cover prior to the 2014 landslide. Evergreen conifer timber harvest over most of the Whitman Bench in the 1940s, with subsequent regrowth of Douglas fir (Pseudotsuga menziesii) trees, created a younger forest, with trees generally less than 20 m tall. On the Paleo landslide scarp, much taller coniferous trees (25–40+ m) were present, because recent logging activities in the late 1980s and early 2000s were restricted in this zone directly upslope of historical landsliding ( Miller and Sias, 1998 ; Everett and Casey, 2014 ). Farther down the slope, in the area of the Hazel landslide, a mix of deciduous and coniferous trees of much shorter height (<15 m) predominated, resulting from (nonplanted) postlandslide regrowth.

Our mapping of features used to infer mobility included the presence and location of evidence of liquefaction such as sand boils and slosh pits, the size and structure of hummocks, and the before- and after-event locations of anthropogenic features such as dislodged pieces of log revetment wall (many with broken steel cables). Sand boils are a classic phenomenon resulting from both cyclic and monotonic (often referred to as “static”) liquefaction (e.g., Seed, 1968 ; Bardet and Kapuskar, 1993 ; Ojha et al., 2001 ) and represent the surface manifestation of eventual post-emplacement pore-water pressure dissipation from depth. They are generated as elevated pore-water pressures eventually move fluid to a free (ground) surface, typically carrying liquefied silt-, sand-, and sometimes gravel-size particles ( Kayen et al., 2004 ) with upwardly flowing water. We readily identified sand boils in the field by their circular, mostly cone-shaped structure and composition of nearly uniform sand-sized particles. Slosh pits are topographic depressions, typically located between hummocks, that contain a mixture of sand and grayish pebbles and cobbles, with a distinctive “bathtub” ring around the margins of the depression. These features likely formed from movement (so-called “sloshing”) of sediment-laden liquid during rapid landslide emplacement. In many cases, sand boils and slosh pit deposits were no longer visible 1 yr following the landslide as a result of ongoing overall erosion at the site resulting from precipitation and bioturbation.

We mapped major scarps, grabens, thrusts, and hummocks (i.e., multisided landslide elements that commonly preserve the uppermost topographic surface) to establish relative senses of motion of the various components of the landslide. The direction, position, and disposition (whether buried, dragged, or fractured) of mobilized trees also provided indications of whether distinct landslide zones experienced extension or compression (or both). We mapped the four major (disrupted) geologic units of the landslide (Q goe , Q gtv , Q gov , Q glv ; see Dragovich et al., 2003 ) by making on-site observations at ∼1400 locations and carefully correlating these observations with known stratigraphy of the area from previous downhole drilling and coring at the site ( Riemer et al., 2015 ; Badger, 2016 ), as well as from scarp exposures. In many cases, thin, disrupted surficial materials concealed the primary units forming individual hummocks; we therefore performed shallow digging and trenching to identify these primary units.

Detailed field geologic mapping and observations form our primary data. The majority of our field investigations took place between April and September 2014 to take advantage of the relatively undisturbed, postfailure geomorphic conditions, with several weeks of follow-up field work over the next 2 years. We also mapped a 450-m-long section through the west distal lobe (see Fig. 2 ) made available via the excavation of the SR530 road bed 2 months following the landslide. Timely mapping on high-resolution (1:1200 scale) orthophotos was crucial, as many features underwent rapid disintegration in subsequent months from both anthropogenic (deposit regrading) and meteorological (rainfall) impacts. For some aspects (e.g., mapping of debris-flow limits), our mapping was aided by photographs and video taken from helicopters by emergency responders during the first 48 h following the landslide.

Geotechnical testing included soil characterization tests (i.e., grain size, Atterberg limits, specific gravity) on samples to determine their provenance and relation to various components of the landslide. We also measured bulk density and water content, and we calculated dry density (when applicable) to obtain typical average values for subsequent mobility analyses. Finally, we conducted a suite of triaxial geotechnical strength tests to investigate potential failure modes of the underlying alluvium sediments forming the North Fork Stillaguamish River floodplain. These tests were performed at densities, saturation levels, and confining stresses commensurate with the field conditions at the time of the landslide. We present additional details of the geotechnical laboratory testing program in Appendix S1 of the Data Repository. 1

We collected sediment samples from key exposures that provided insights into possible controls on landslide mobility. These included debris-flow deposits (Q df ) sampled at the distal edge of the landslide shortly (weeks) following emplacement, sand boil samples located throughout the hummock field, overridden floodplain alluvial sand (Q a , as identified and mapped herein) from locations temporarily exposed as the new channel of the North Fork Stillaguamish River cut through the landslide deposit, and recent, post-2014, alluvium deposited along the North Fork Stillaguamish River following the landslide (as a proxy for the deposits forming the floodplain prior to being overrun by the landslide).

Multiple high-resolution aerial LiDAR data sets acquired before (July 2013) and after (26 March 2014 and 6 April 2014) the 2014 landslide allowed us to analyze three important characteristics: (1) the thickness of the slide deposits, (2) the three-dimensional (3-D) geometry of the slide surface, and (3) the heights of the preslide trees (see Appendix S2 in the Data Repository [ footnote 1 ]).

We calculated the spatial distribution of slide-deposit thickness by first constructing bare-earth digital elevation models (DEMs) for each LiDAR point-cloud data set. This entailed using last-return laser data and performing minor manual editing to remove vegetation and anthropogenic structures from each DEM. We then compared independently georeferenced pre- and postslide pairs of these DEMs to determine elevation differences. Vertical differences between unchanged areas of the models (i.e., away from the landslide) were less than 20 cm for 90% of the data (and less than 100 cm for 99.9% of the data), thereby allowing a spatially robust volumetric comparison of areas affected by the landslide. Areas on the North Fork Stillaguamish River valley plain that increased in elevation following the 2014 landslide indicated slide-deposit thickness (assuming only minor scour and/or compaction of the subsurface). In contrast, elevation changes on the hillslope may have resulted from rearrangement of mass by sliding and were not indicative of deposit thickness. For these areas, we constructed surfaces of the inferred slide plane at depth based on borehole and other data (see next section and Appendix S2 of the Data Repository [ footnote 1 ]).

Combined with borehole data showing 2014 failure surface elevations at five locations within the landslide deposit ( Badger, 2016 ), we used the LiDAR data to reconstruct inferred failure surfaces of the landslide. Using this information in a fully 3-D combination LiDAR- and CAD-based software platform (Maptek I-Site Studio v. 5.1), we constructed both failure and ground surfaces, and we subsequently calculated volumes of both the source and deposit regions. These included the identification and calculation of separate components of the source and deposit areas (e.g., the volume of the Paleo landslide source and the Hazel landslide source). Details of our volume calculations are included in Appendix S2 of the Data Repository ( footnote 1 ).

We used particle-tracking methods focused on readily visible “tall” (>20 m in height) trees in both aerial photographs and LiDAR data to aid the quantification of mobility between source and deposit zones of the landslide. Prior to sliding, taller Douglas fir trees (20–40+ m) were restricted to parts of the landslide source areas, mostly in the upper half of the slope. Other areas lower on the slope were covered with much shorter (typically <15 m in height) mixed deciduous and coniferous forest. After the 2014 event, trees were displaced and distributed throughout various parts of the slide deposit, with most trees toppled over and having subhorizontal orientations. We determined preslide tree heights using the 2013 LiDAR point cloud, where the difference in elevation between first- and last-point LiDAR returns provided us with estimates of individual tree height. Then, using a 14 April 2014 postslide orthophoto, we located and measured the lengths of toppled, subhorizontal “complete” conifer trees, distinguished by visible root systems and pointed, unbroken tips. Numerous other downed “partial” trees existed on the deposit (as noted by Keaton et al., 2014 ; Wartman et al., 2016 ), but we focused on complete trees.

Our geologic and structural cross section through the landslide ( Fig. 8 ) used data from 12 geotechnical boreholes ( Fig. 4 ; Badger, 2015 , 2016 ) along with our field mapping to portray the disturbed stratigraphy within the landslide deposits. We used a turning point located where major block slices transition into hummocks to present the entire section of the landslide, from source zone to distal deposit. This view provides the subsurface details of the various slide components (i.e., major deposit zones, failure surfaces, etc.), which were subsequently used to interpret the landslide sequence (see next section).

The deposit thickness and sand boil map ( Fig. 6 ) shows the locations of 107 sand boil sites (composed of one or more of the 379 individual sand boils we mapped; see Data Sheet S2 of the Data Repository for location data of sand boils [ footnote 1 ]) overlain on an elevation change map, where positive values on the valley floor approximate 2014 slide deposit thickness. Most sand boil sites were located south of the former footprint of the northern margin of the North Fork Stillaguamish River, with only two locations mapped north of this boundary. Sand boils were typically several decimeters in diameter and several to ten centimeters tall, although in some cases, they reached over 1 m in diameter and several decimeters tall ( Fig. 7 ). We used sand boils as indicators of liquefaction at depth. We found sand boils in both subaerial and subaqueous environments, indicating that pore-pressure dissipation occurred shortly after emplacement. Some sand boils were located within slosh pit features ( Fig. 6 ); sediment rings were higher on the distal sides of these slosh pits, indicating rapid deceleration.

Our mobility features map ( Fig. 5 ) portrays the tracking of several groups of objects from their preslide to postslide locations; these objects include complete trees, pieces of forest floor, and pieces of a cut-log revetment wall originally located on the north side of the North Fork Stillaguamish River. The locations of these features could be readily identified using pre- and postslide aerial photos, LiDAR-generated tree-height maps, and field observations (see “Methods” section). Although we did not perform a one-to-one correlation between each of these features, patterns of movement are evident. For example, tall trees originally standing on the scarp and bench of the Paleo slide as well as a small section of the Whitman Bench are readily distinguished in their fallen locations on various parts of the landslide deposit (see color-coded tree heights in Figs. 5A and 5B ; data provided in Data Sheet S3 of the Data Repository [ footnote 1 ]). Here, we also show areas of disaggregated and stretched forest floor ( Fig. 5B ) on the surface of the landslide deposit, along with pre- and postslide locations of parts of the log revetment structure.

Our geologic map ( Fig. 4 ) depicts broad swaths of the main stratigraphic units that make up the source and deposit areas. These consist of the four predominant glacial deposit units (Q goe , Q gtv , Q gov , Q glv ; Dragovich et al., 2003 ) as well as several units that are specific to the landslide’s depositional structure. Units we mapped and informally refer to here include areas where hummock composition is highly mixed (undifferentiated hummocks, Q gund ), the large colluvial infill downslope of the landslide headscarp (Q c ), and debris-flow deposits located at the distal (Q df ) and lateral (Q dfs ) margins of the landslide. Our original mapping data consisting of lithologic observations are presented in Data Sheet S1 of the Data Repository (see footnote 1 ).

Our structure map ( Fig. 3 ) presents major scarps, grabens, thrusts, and blocks throughout the landslide and identifies ∼900 hummocks in the central to distal end of the deposit. Individual hummocks were identified using weighted and smoothed slope curvatures of the 2014 LiDAR-based DEM at three scales (3–10 m, 25–50 m, and 50–100 m) to isolate flat hummock tops. Geomorphic features transition from large downdropped blocks to laterally continuous block “slices,” hummock fields, and smaller disaggregated hummocks in progression from north to south through the deposit. Nearly all structural features show evidence of extension, with a few notable exceptions described in the “Analysis and Interpretation of Landslide Progression” section.

Landslide mobility is commonly quantified using travel paths of particular features. For example, tracking where an anthropogenic (e.g., piece of revetment wall) or natural feature (e.g., tree) was located before and after a landslide can provide a direct indication of surface displacement. At the Oso landslide, strong evidence for travel paths is provided by the three major components of the landslide source area, each with its own stratigraphic package (Hazel, Paleo, and Whitman; Figs. 2 and 4 ). The Hazel source component was located just north of the North Fork Stillaguamish River and consisted of debris from previous landslide failures. The Paleo source component was located just upslope (north) of the Hazel component and consisted of the remnants of a downdropped block that (prehistorically) failed from the Whitman Bench (see “Study Area Background and Landslide History near Oso” section). Finally, the Whitman source component consisted of the part of the Oso landslide that failed within intact geologic units forming the Whitman Bench. The Whitman component also produced several secondary failures that, although not directly involved in the Oso landslide’s mobility, are important for understanding the deposit geometry. In this section, we use our mapped lithologic relations ( Figs. 4 and 8 ), along with observations of displaced trees, forest floor, and revetment wall ( Fig. 5 ), to identify, interpret, and link the various source ( Fig. 9A ) and deposit ( Fig. 9B ) areas of the landslide (Hazel, Paleo, Whitman). Using these linkages, we can then interpret the landslide sequence (presented in a subsequent section), which is in turn necessary for understanding the timing and mobility of the landslide.

Finally, additional confirmation of a dual Hazel and Paleo source for the deposits south of the C-E zone is possible through volume-balance calculations. Using 3-D reconstruction of pre- and postslide topographic surfaces and construction of the landslide sliding surface (see “Methods” section), we estimated the combined volumes of the 2014 Hazel and Paleo deposits south of the C-E zone (including debris flow) to be ∼2,875,000 m 3 . The volumes of the Hazel and Paleo component source areas were 1,930,000 m 3 and 766,000 m 3 , respectively, making for a combined volume of 2,696,000 m 3 and resulting in a volumetric expansion ratio of ∼7% ( Table 1 ). This somewhat low expansion ratio (compared to rock avalanches; e.g., Nicoletti and Sorriso-Valvo, 1991 ) suggests that the majority of the interparticle motion was restricted to the basal slide surface, with large extension between hummocks accounted for in the LiDAR-based volume calculation. Any additional source area volume that might have been part of the Hazel-Paleo failures would result in an even smaller expansion ratio, and assuming that all deposits south of the C-E zone came only from the Hazel source area results in an unrealistically large (∼50%) expansion ratio. Thus, the volumetric calculations reinforce the interpretation that the deposits south of the C-E zone originated from the combined Hazel and Paleo source areas ( Figs. 8 and 9 ).

Our mapping of the distribution of postsliding revetment logs and vegetation ( Fig. 5 ) serves to further distinguish the provenance of the hummock deposits south of the C-E zone. We mapped revetment logs with cables in numerous locations on the southern border of the hummock field (i.e., primarily among and distal to the undifferentiated hummocks). These logs originally formed part of the erosion-protection revetment structure located at the toe of the pre-2014 Hazel landslide deposit. As such, these dispersed logs denote a Hazel component origin, consistent with our interpretation for the source of the undifferentiated hummocks. Transported remnants of stretched forest floor ( Fig. 5B ), some with tall (20+ m) trees still intact, also occupy the tops of many hummocks south of the C-E zone and indicate stretching of the deposit during landslide movement. Although a few trees south of the C-E zone remained somewhat upright after sliding, most fell or were snapped. We identified numerous complete trees (see “Data and Observations” section) with heights of 20–40+ m lying within the stratified (inferred Paleo) hummock zone ( Fig. 5B ). Trees of this height were prevalent in the Paleo bench area but virtually nonexistent in the Hazel slide area prior to the 2014 event, thereby helping to distinguish Paleo from Hazel components in the deposit south of the C-E zone ( Fig. 9B ).

In contrast, the hummocks just north of the inferred Hazel deposits (but still south of the C-E zone; Fig. 3 ) are primarily stratigraphically conformable (i.e., in sequence of Q goe –Q gtv –Q gov –Q glv ), with the deepest materials from the originally horizontally layered glacial stratigraphy sequence being transported the farthest to the south ( Fig. 4 ). Some unconformities exist locally in the hummock field (e.g., Q gtv is missing in the far western lobe, and Q gov is missing in the central distal end); however, these can be explained by source-area stratigraphic relations and material behavior during transport. Large swaths of intact Q gtv are mostly absent on the western side of the Whitman Bench source area, and thus its absence in parts of the western hummock field is expected. Elsewhere, in the central part of the deposit south of the C-E zone, blocks of stronger, rock-like Q gtv may have tumbled and overridden weaker (i.e., uncemented sand) units such as Q gov during transport, providing an explanation for the absence of Q gov deposits here. This behavior is also visible on the east side of the Whitman component of the 2014 landslide (discussed subsequently), where a large swath of Q gtv blocks overrode (presumed) areas of Q gov ( Fig. 4 , near boreholes EB-07 and EB-18). Overall, the structured nature and distinguishable lithology of the hummocks south of the C-E Zone strongly suggest that they came from a large, and more intact, source area north of the (prelandslide) river. We infer that the Paleo bench area was the source for this part of the 2014 deposit (as opposed to being composed of a relic, undisturbed section of sediments). Importantly, we found no structural evidence of compressional or extensional features separating the primarily stratified (Paleo) from the undifferentiated (Hazel) hummock zones that might indicate separate failure times for these components. Our observations therefore indicate synchronous failure of the combined Hazel-Paleo source for the hummock deposits south of the C-E zone.

Our mapping south of the C-E zone revealed a southward progression of hummocks composed mostly of lithologically distinguishable materials (i.e., recessional outwash [Q goe ], till [Q gtv ], etc.). Toward the southern end of the hummock field, but north of the debris-flow deposits ( Fig. 4 ), we identified a swath of chaotically mixed smaller hummocks (i.e., till, advance outwash, and advance glaciolacustrine deposits all in close, but seemingly random, proximity) juxtaposed against larger more monolithologic hummocks to the north. This mixed zone is labeled “undifferentiated hummocks” (Q gund ) in Figure 4 , and we interpret this zone as originating from the Hazel landslide area north of the North Fork Stillaguamish River. Repeated cycles of landsliding at the Hazel site during its 80+ yr history prior to the 2014 landslide likely mixed the glacial units into a mélange, as reflected in the composition of the undifferentiated hummocks. These hummocks, along with the distal debris-flow deposit, compose the Hazel component deposit.

Using our geologic mapping of the 2014 deposits, we first examined the distal areas of the Oso landslide deposit south of the North Fork Stillaguamish River, and in particular, south of the compression-extension (C-E) zone ( Fig. 3 ) and north of the distal debris-flow deposits ( Fig. 4 ). The C-E zone forms an important structural boundary between two of the deposit areas and is discussed in more detail in the subsequent section.

Most of the overall Oso landslide deposit exhibits extensional features ( Fig. 3 ). However, the northern boundary of the Paleo deposit differs markedly from other features within the deposit. This internal feature represents the contact between deposits from the Hazel-Paleo source area and deposits from the Whitman source area (described subsequently). This contact is readily identifiable by repeated stratigraphy upslope and downslope of the contact zone ( Fig. 4 ) and by vegetation patterns, where upslope Q glv deposits (inferred Whitman) with few trees are (unconformably) in contact with downslope Q goe deposits (inferred Paleo) covered with many toppled trees and pieces of stretched forest floor ( Fig. 5B ). The contact between these two masses is distinctively preserved south of the current North Fork Stillaguamish River ( Fig. 10 ); we made additional detailed field observations to identify interactions in this zone. Whereas we found localized evidence of compression along the length of the contact (isolated snapped and/or buried trees), we also found ample evidence of extension (grabens, trees toppled back onto the Whitman deposit, and trees with drag marks). Grabens are present along ∼60% of the contact between the Whitman and Paleo deposits ( Fig. 10 ; as calculated over the length of contact of the Whitman Q glv deposit with the Paleo Q goe deposit; Fig. 4 ). With the exception of one distinct region of compression on the east containing a high concentration of snapped and buried trees and minor thrusts in the absence of any grabens, few other areas along the contact zone contain exclusively compressional features. Widespread compressional features would be expected if the massive upslope Whitman component impacted a static Hazel-Paleo component. The presence of toppled trees (rooted entirely in Paleo component materials) on top of the Whitman component also indicates that the two components were moving simultaneously. We therefore denote this area of juxtaposed compression and extension as the compression-extension (C-E) zone and interpret its kinematics to infer that the distal edge of the Whitman component and the proximal edge of the Paleo component, although distinct entities, were traveling at similar rates and in close proximity to one another.

The total source volume of the Whitman component (6,311,000 m 3 ) was considerably larger than the combined masses of the Hazel-Paleo slide ( Fig. 8 ; Table 1 ), and it failed along a slip surface with maximum depth of ∼100 m. The overall shape of the failure surface (as revealed by geotechnical borings and other field observations; see Appendix S2 of the Data Repository [ footnote 1 ]) is arcuate near the headscarp, but translational at a shallow dip over the majority (60%, 330 m) of its length. This shape may be the result of the anisotropy of the dominant (and horizontally bedded) Q glv sediments, which formed ∼87% (480 m) of the deeper failure surface length (∼550 m; Fig. 8 ). Anisotropy in clays is known to affect soil shear strength (e.g., Duncan and Seed, 1966 ; Kirkgard and Lade, 1991 ) and failure plane shape (e.g., Lo, 1966 ; Badger and D’Ignazio, 2018 ).

Given the C-E zone contact, the surface expressions of source and deposit areas for the Whitman component of the Oso landslide are more distinct than the Hazel and Paleo components. The previously unfailed Whitman source area ( Fig. 9A ) is characterized by a section of intact and mostly horizontally conformable glacial deposits (i.e., layered Q goe –Q gtv –Q gov –Q glv from top to bottom) behind the crest of the Paleo landslide headscarp ( Fig. 2 ) and is part of the Whitman Bench topography. The 2014 Whitman deposit, located immediately below (and to the south of) the 2014 Oso headscarp ( Fig. 9B ), is characterized by large extensional features (i.e., blocks, slices) and exposed scarps and, on flatter ground near the C-E zone, large (10–15-m-tall) hummocks. Although some of the Whitman component traveled out onto the North Fork Stillaguamish River valley floor, most of the mass remained on the slope north of the river valley. The uppermost section of the Whitman component is a block ( Fig. 3 ) composed primarily of recessional outwash that was back-rotated ∼10°–15° and covered with extensive downed trees with tips pointing northward, suggesting rapid downslope acceleration. The interior of the block contains extensional features with transverse scarps and fractures. Downslope of this block, the mass extended and developed multiple slices separated by listric (internal) faults ( Figs. 3 and 8 ), which formed as the slide translated along a slightly inclined (6°) basal failure surface ( Fig. 8 ). These slices created scarps up to 35 m tall, with some (e.g., Fan Lake scarp; Figs. 3 and 8 ) nearly exposing the Oso basal failure surface. The slice detachment from the Fan Lake scarp (and possible partial collision into the Paleo component ahead of it) may have caused some of the compressional features located on the west side of the C-E zone ( Fig. 10 ). Original stratigraphic relations are mostly preserved in the translated Whitman deposit, and at several locations close to the new alignment of the North Fork Stillaguamish River our mapping ( Fig. 4 ) revealed displaced, but nearly horizontal, stratigraphic contacts for till through glaciolacustrine units (i.e., Q gtv over Q gov over Q glv ). Relatively intact blocks of till (which is locally extensive on the east side of the exposed Whitman Bench scarp) were scattered and transported nearly 500 m along the eastern half of the Whitman component ( Fig. 4 ).

The energy associated with the colluvial infilling (sliding and/or falling material, inclusive of the till falls) could have been considerable (estimated to be ∼1150–1950 GJ if the entire mass, with an assumed intact unit weight of ∼2040 kg/m 3 , experienced a free fall of 90 m [upper value] or if it experienced Coulomb frictional sliding along a roughly 50° backslope with a length of 100 m and a friction angle of 30° [lower value]). Although this infilling may not have contributed directly to the mobility of the 2014 landslide, it may assist interpretation of the recorded 2014 seismic signals from the landslide (see subsequent “Discussion” section). Characteristics of these colluvial materials also provide insight into the composition of the earlier Paleo source area prior to its failure in 2014. For example, the predominantly loose, sandy, forest floor covering recessional outwash deposits (Q goe ) found in hummocks at the back edge of the 2014 Paleo landslide deposit contains seemingly random blocks of till (Q gtv ) and glaciolacustrine clay (Q glv ; mapping locations are provided in Data Sheet S1 of the Data Repository [ footnote 1 ]). These out-of-stratigraphic-sequence blocks may be the result of similar colluvial infilling in the graben of the prehistoric Paleo landslide.

The other secondary component of the Oso landslide consisted of an ∼1.1 million m 3 section ( Table 1 ) of the intact Whitman Bench glacial deposits that subsequently slid or toppled into the graben formed behind the downdropped back-rotated Whitman block. This deposit (denoted as colluvium, Q c ; Figs. 4 and 8 ) is composed of juxtaposed outwash and till deposits (i.e., Q goe , Q gtv , Q gov ), with modern roots, wood, and other organics (as identified in geotechnical boring EB-04si-15 at 26–38 m depth and within 4 m of the Whitman slide failure surface; Badger, 2016 ). The presence of recently buried organics at this location suggests that the primary Whitman failure surface intersected the Whitman Bench well in front (to the south) of the final location of the escarpment (see inferred failure surface shown in Fig. 8 ). The upper part of the colluvial infill consisted of a series of at least three distinct so-called “till” fall events (totaling ∼41,000 m 3 ; Table 1 ), with the largest event encompassing ∼75% of the total till-fall volume. The largest of these falls, composed of competent blocks of well-cemented glacial till, partially disaggregated upon impact and subsequently spread and buried part of the upper Whitman slide block, splintering back-thrown, 40-cm-diameter trees that had been previously growing on recessional outwash on the top of the Whitman Bench.

Two additional aspects of the Oso landslide’s overall structure are notable for their evidence of landslide timing and structure, despite their lack of direct correspondence to overall landslide mobility. These secondary failures were part of the Whitman component, but they occurred subsequent to the major motions of the Hazel, Paleo, and Whitman components. The first consisted of a debris flow along a 400 m length on the east side of the landslide source area ( Fig. 4 , Q dfs ) that formed soon (<1 h) after the main landslide (see video from search and rescue helicopters, available at https://www.youtube.com/watch?v = UUFByAwcGs0, accessed 23 May 2017). This debris flow mobilized glaciolacustrine clay (Q glv ), likely as a result of groundwater capture from the neighboring drainage basin to the east ( Keaton et al., 2014 ). Notably, this is the only major area of the Whitman component that mobilized into a flow.

Here, we present our interpretation of the sequence of events of the 2014 Oso landslide. The sequence is important not only for understanding the failure mechanics and kinematics of the landslide, but also because it points to factors that likely enhanced the landslide’s mobility. We base this sequence primarily on our field observations and analysis of the source and deposit components, as presented in previous sections. We organize the sequence into three subsections of motion: (1) slope preconditioning and landslide triggering, (2) initial motion, and (3) landslide transport. The final subsection is further divided into two parts, describing leading-edge debris-flow formation and landslide runout. When discussing landslide runout, we provide detailed information on our proposed mobility mechanism (valley bottom liquefaction), which sets the stage for our subsequent quantitative analyses of this mechanism. Each of the sequence stages references various components of the landslide (Hazel, Paleo, Whitman) as previously described.

Finally, it has been hypothesized that elevated pore-water pressures in the subsurface, induced by infiltration from significant rainfall in the weeks or years leading up to failure ( Henn et al., 2015 ), may have triggered sliding ( Iverson et al., 2015 ; Stark et al., 2017 ; Wartman et al., 2016 ). On the other hand, if a progressive strength-loss mechanism triggered initiation ( Badger and D’Ignazio, 2018 ), elevated pore pressures would not necessarily have been needed. It is possible that some type of hybrid failure mechanism controlled initiation, such as progressive failure of lacustrine clay in the Whitman component combined with pore-pressure–induced failure (with possible loading from the upslope Whitman component) in the Hazel-Paleo component. Although we do not specifically address initial triggering in our analyses, it is clear that the 2014 Oso landslide occurred on a slope that was preconditioned by prior landsliding (i.e., the Hazel and Paleo components) for future instability.

The site of the Oso landslide consisted of a 180-m-tall terrace of glacially derived sediments on the north side of a valley that had been previously subject to extensive fluvial incision at its base. Although the toe of the 2014 failure surface was above river level at the time of failure ( Fig. 8 ), toe erosion did establish the overall geometric configuration of the slope. Moreover, the stress condition of the slope also likely played a role in its failure, with large slope height (and resultant higher stress state) leading to lower overall stability ( Badger and D’Ignazio, 2018 ).

Piecing together the sequence of postfailure landslide mobility for the Oso landslide requires some understanding of the reasons for its initiation. Although we deliberately focused on the Oso landslide’s mobility in this study rather than on its initiation, triggering factors can shed light on aspects of mobility. For example, hydrologic conditions are generally important, with fluidization of water-saturated sediment leading to larger mobility (e.g., Legros, 2002 ; Okura et al., 2002 ; Iverson et al., 2011 ). Landslide volume and topography (e.g., Corominas, 1996 ), as well as failure history, can also affect landslide mobility. For example, the La Conchita, California, landslides in 1995 and 2005 occurred on a slope with evidence of prehistoric landsliding ( Jibson, 2005 ). To this end, the history of landsliding at the Oso site helps to place the 2014 slide’s prefailure state into context.

Given these relations, our interpreted sequence begins with the Hazel-Paleo components sliding (and potentially dropping slightly) from the toe of the basal slide surface located ∼10 m above the existing ground surface at the base of the slope ( Fig. 8 , section distance 530 m to 560 m). As the Hazel-Paleo component moved out into the valley, it pushed and entrained the North Fork Stillaguamish River immediately in front of it, which was transported across the valley (see subsequent section). The Whitman component, now partially exposed by the backscarp of the newly evacuated Paleo component ( Fig. 8 ), followed closely behind the still-moving Hazel-Paleo mass. As the Whitman component moved southeastward, it rounded a resistant knob (known as the “Pyramid”; Fig. 3 ) composed of downdropped deposits displaced from prior prehistoric landsliding and collocated with the Devils Mountain fault trace ( Johnson et al., 2001 ; Dragovich et al., 2003 ). All three components then spread into the unconfined space of the North Fork Stillaguamish River valley floor.

Although our field evidence indicates that both components were moving closely in time, a precise failure mode is not directly apparent. Geological and structural evidence north of the C-E zone indicates that the Hazel-Paleo component pulled away from the Whitman component by a sufficient distance to allow exposure and subsequent extension of the Q gov and Q glv parts of the Whitman mass ( Figs. 4 and 8 ). If the Whitman component failed first (progressive failure in Q glv sediments), with the Hazel-Paleo component sliding away in response, a large separation of the two masses concurrent with jostling in the C-E zone might be less likely. Thus, our mapping suggests a more retrogressive failure mechanism, with the Hazel-Paleo component moving prior to the Whitman component. Regardless of whether initial motion was progressive or retrogressive, all source components failed closely in time, thereby generating considerable momentum. This has important ramifications for mobility in that the Oso landslide moved out onto the river valley in two closely spaced components, but as one mass.

Regardless of the exact mechanism for initiation (precipitation-induced pore pressure changes, progressive strength loss, or other), our analysis of field evidence (see “Relation between Landslide Components” section) indicates that the Hazel and Paleo components failed together as one mass and that they likely initiated just shortly before the Whitman component began to slide. The short time gap in initiation of these two components is evidenced by both compression and extension in the C-E zone ( Fig. 10 ). If a static Hazel-Paleo component was impacted by later failure of the Whitman component, the C-E zone would show predominantly collisional or compressional features. Alternatively, a postfailure static Whitman deposit with secondary motion of the Hazel-Paleo component in front of it would result in solely extensional features in the C-E zone. The presence of both features in this zone indicates jostling between the components and that they were in motion simultaneously.

Several events, including the formation of a debris-flow front and transport of the slide mass across the North Fork Stillaguamish River valley, occurred quickly following initial motion of the landslide. Other investigators have referenced these events to seismic records and eye-witness accounts; we compare our findings to these previous versions in the subsequent “Discussion” section. Here, we use and discuss our field observations to support our inferred sequence for landslide transport and mobility.

River entrainment and debris-flow formation, runout, and reflection. Immediately after the 2014 landslide, the most distal parts of the deposit contained fully liquefied materials with a perimeter snout of logs and debris, characteristic of a debris flow. The debris flow represented the leading edge of the landslide mass; as it swept across the floodplain, it entrained nearly all surface features (trees, cars, and structures, including victims) that previously existed on the south side of the river. Pieces of the log revetment formerly located on the north side of the river were also found within some parts of the debris-flow deposit (Fig. 5). Many of these items were transported to the outer margins of the deposit (Fig. 4D), as evidenced by the locations of structural debris and victim recovery (Quistorf, 2014). In the first weeks after the event, we observed liquefied pools of sediment and ponded water at the distal end of the landslide deposit formed from debris-flow emplacement; these were a significant hindrance to rescue workers (Fig. 4D).

Field samples we collected several weeks after the landslide at the distal edges of the debris-flow deposit revealed an average bulk density of 1950 kg/m3 (Table S1 [footnote 1]). With a density roughly twice that of water, the rapidly moving debris-flow slurry had sufficient momentum to sweep houses from their foundations. However, few indications of subsurface entrainment or scour within the Steelhead Haven neighborhood were found during our or other previous investigations (Iverson et al., 2015; Wartman et al., 2016). Our mapping did identify two sections of the SR530 road bed, with a combined length of ∼150 m, that were disrupted by the landslide (Fig. 5B), but they appear to have been part of engineered embankments built 3–4 m above the adjacent topography.

The debris flow likely formed as the toe of the combined Hazel-Paleo component crossed the relatively flat, 200-m-wide area just north of the North Fork Stillaguamish River, struck the 2006 log revetment along the north edge of the North Fork Stillaguamish River, and entrained the river. Prior to the 2014 event, an ∼10-m-thick wedge of colluvium (tapering from 21 m to 1 m in thickness; Fig. 8), derived from 2006 Hazel slide debris, had accumulated behind the roughly 4-m-tall log revetment (Keaton et al., 2014; Wartman et al., 2016). The leading edge of the 2014 slide may have mobilized this flat-lying debris north of the revetment as well. However, given that the failure surface toe emerged from the slope above this older colluvial debris (Fig. 8), the Hazel-Paleo component may have primarily overrun this debris before impacting and entraining the log revetment. Collapse during failure of loose Hazel component materials on the slope may have aided debris-flow formation at the front of the moving mass, although field evidence of hillslope liquefaction is minimal. A significant source of water for debris-flow mobility likely came from entrainment of the North Fork Stillaguamish River, with the remaining being in the previously failed debris. Based on time-series comparison of measured and correlated flows in the river at the time of the 2014 landslide (see Appendix S2 in the Data Repository [footnote 1]), we estimate the entrained river volume to be ∼78,000 m3 or roughly 18% of the debris-flow deposit (424,000 m3; Table 1).

The debris flow was unconfined as it traveled onto the flat North Fork Stillaguamish River valley, and it moved in tandem with the mass of the combined Hazel-Paleo component immediately behind it. Our mapping shows that the debris flow reflected off the southern central wall of the valley, flowed back toward the north, and onlapped onto the hummock field of the Hazel component of the deposit (Fig. 11). This reflection indicates that the debris flow remained close to and in front of the Hazel-Paleo component as the components traveled across the valley. Where the debris flow was not reflected (i.e., in the eastern and western lobes), the debris thinned as it traveled farther, as evidenced by measurements of mud onlap deposits on the base of trees (Fig. 11A). These thinner deposits commonly contained small hummocks (typically <1 m high) of various glacial units that progressively decreased in size with farther travel distance. Average distal tree onlap measurements provide a reasonable indication of the distal deposit thickness (0.8 m and 1.2 m in the west and east lobes, respectively; see Appendix S2 of the Data Repository [footnote 1]), and they likewise help to constrain the volume (424,000 m3) of the debris flow. Although the debris-flow component of the event caused significant damage, it only represented ∼4% of the total landslide deposit volume (Table 1).

Landslide runout and extension via valley bottom liquefaction. After initial motion and subsequent entrainment of the river, the combined Hazel-Paleo component, along with the southern part of the Whitman component, traveled rapidly across the North Fork Stillaguamish River floodplain immediately south of the slide (Fig. 9). The deposit from this part of the landslide mass was primarily a field of hummocks, characteristic of a debris avalanche. The hummocks formed when the mass extended and spread in multiple directions into the accommodation space offered by the open, relatively flat North Fork Stillaguamish River valley bottom. Hummock materials were not liquefied, and most of their cores consisted of one stratigraphic unit (i.e., Q goe , Q gtv , etc.), except in distal smaller hummocks, where mixed materials were common. Upright ferns and forest floor also remained intact on hummock tops (Fig. 5B). Proximal hummocks were typically ridgelike, with longer axes oriented perpendicular to the direction of motion. Many distal hummocks displayed nearly equidimensional geometries, indicating that extension occurred similarly in both travel-parallel and travel-perpendicular directions. In some cases, hummocks themselves underwent internal extension with jigsaw-puzzle fabric and vertical cracks extending tens of decimeters downward into the hummocks (Fig. 4C). These features are common in other debris-avalanche hummocks (e.g., Roberti et al., 2017).

As a feature of extensional kinematics, hummocks typically indicate the presence of a weak base underlying the landslide mass (Paguican et al., 2014). The relatively flat North Fork Stillaguamish River valley bottom beneath the Steelhead Haven neighborhood, subsequently buried by hummocks (Fig. 12), consisted of alluvial sediments (primarily silts, sands, and gravels) ranging from <1 to 13 m in thickness (Fiske, 2014; see also Table S5 [footnote 1]). Given the anomalously high precipitation and resultant soil moisture during March 2014 (Henn et al., 2015), near-surface alluvium was likely saturated or nearly saturated with shallow groundwater near the ground surface. Geotechnical borehole drilling conducted 1 month following the landslide (Fiske, 2014) encountered groundwater within 1–2 m of the preslide ground surface over much of the distal edge of the deposit, even after a temporary lake that had formed by the landslide blocking the North Fork Stillaguamish River had drained.

Several lines of field evidence indicate that liquefaction of the North Fork Stillaguamish River valley alluvium was the likely source for this weak base, and for the high mobility of the slide mass. The first is the occurrence of the hummocks themselves. Lateral deformation resulting from liquefaction often results in hummock-like features (e.g., Hansen, 1965; Harty and Lowe, 1995; Davies, 2003), and thus liquefaction of the underlying alluvium offers a compelling explanation for their formation. The second line of evidence is a strong spatial correlation between the distribution of hummocks and the location of the flat, pre-2014 North Fork Stillaguamish River alluvial valley. Our structural map (Fig. 3) shows that nearly ubiquitous extension of hummocks exists only south of the earlier (2003) location of the North Fork Stillaguamish River (near the present-day location of Fan Lake; Fig. 8). On the other hand, the area north of Fan Lake (within the northern part of the Whitman deposit) is composed mainly of large, linear, slice-like blocks with their long axes oriented perpendicular to the direction of landslide motion. The transition from larger scarps in the north to hummocks in the south occurs at the location of the start of the preslide floodplain and indicates that basal liquefaction likely occurred to the south of this transition point.

The third key piece of evidence for basal liquefaction-induced mobility comes from observation and identification of hundreds of liquefaction features (i.e., sand boils; Figs. 6 and 7) throughout the deposit zone south of Steelhead Lake (Fig. 3). Given the occurrence of sand boils within many slosh pit features, slosh pits may also be liquefaction-related features, possibly formed when underlying alluvium was expelled or incorporated into these depressions. Sand boils and slosh pits are primarily present south of the recent (2013) North Fork Stillaguamish River (Fig. 6), with the vast majority located between hummock mounds where thinner landslide deposits allowed a shorter pathway for pore-pressure dissipation from the liquefied base (Fig. 13). The formation of sand boils through and on top of “backwashed” distal debris-flow deposits (Figs. 6 and 13) indicates that the sand boils were a late-stage depressurization response of the underlying alluvium. Our geotechnical characterization of sampled alluvium and sand boils from the field (Fig. 14) revealed pronounced similarities in their grain-size distributions; the alluvium underlying the landslide deposits (gray shade in Fig. 14) appears to have been the source of the sand boils (red lines in Fig. 14). Of the other geologic units forming the slide mass (Q goe , Q gtv , Q gov , Q glv ), only Q goe has a similar grain-size distribution to that of sand boils; however, it is coarser grained than the sand boils, indicating a poorer match for the provenance of the sand boils.