Interseismic Lithospheric Response of the Southern End of The

INTERSEISMIC LITHOSPHERIC RESPONSE OF THE SOUTHERN END OF THE

CASCADIA SUBDUCTION ZONE SINCE THE 1992 CAPE MENDOCINO M 7.1

EARTHQUAKE

Jessica Vermeer

A Thesis Presented to

The Faculty of Humboldt State University

In Partial Fulfillment of the Requirements for the Degree

Master of Science in Environmental Systems: Geology

Committee Membership

Dr. Mark Hemphill-Haley, Committee Chair

Dr. Raymond “Bud” Burke, Committee Member

Dr. Robert Rasmussen, Committee Member

Dr. Rick Zechman, Graduate Coordinator

May 2016 ABSTRACT

INTERSEISMIC LITHOSPHERIC RESPONSE OF THE SOUTHERN END OF THE CASCADIA SUBDUCTION ZONE SINCE THE 1992 CAPE MENDOCINO M 7.1 EARTHQUAKE

Jessica Vermeer

The Cascadia subduction zone (CSZ) in the Pacific Northwest, where the Gorda

plate is subducting beneath the North American plate, may be capable of producing M 9

earthquakes. At its southern end, the CSZ terminates at the Mendocino triple junction in

northern California, a region of frequent seismic activity. Unique among this seismicity

was the 1992 M 7.1 Cape Mendocino earthquake, which caused up to 1.4 m of measured

coseismic uplift and may have been a segmented rupture of the southern end of the CSZ.

The coseismic deformation was measured using Vertical Extent of Mortality of intertidal organisms, as well as a first order NGS leveling survey. Using static GPS relocation of leveling benchmarks and the position of intertidal organisms I measured vertical crustal deformation over the 23 years since the 1992 event. If this earthquake had occurred on

the megathrust interface between the Gorda and North American plates, I expected to see

10 – 20 cm of subsidence near the peak of coseismic uplift. However, this earthquake

could have occurred along a subsidiary fault within the accretionary wedge or upper

plate, in which case the deformation rate should be much lower. Benchmark relocation

and intertidal organism relocation yield maximum vertical deformation of 1 mm/yr.

These low interseismic deformation rates measurements indicate that the 1992 M 7.1 ii Cape Mendocino earthquake was not rupture of the subduction zone interface, but likely occurred within the upper plate accretionary complex on a subsidiary fault.

iii ACKNOWLEDGEMENTS

First I need to thank my committee, Mark, Bud and Bob, for their support, input and inspiration. Huge thank you to Angela Jayko, USGS retired. She was extremely helpful and went through the trouble of digging up a whole set of photos and notes from the 1992 intertidal organism survey, which I used extensively and made the non-GPS portion of this study possible. Gary Carver, Bill Hammond (UNR), and Pat McCrory

(USGS) all gave me input and suggestions that were invaluable. Meghan Miller and Jim

Normandeau from UNAVCO enabled the GPS equipment loan on extremely short notice once I realized the RTK was not sufficient. I received financial support from the HSU

McCrone Graduate Fellowship and a GSA student research grant. Many friends helped me with fieldwork and I couldn’t have completed it without them: Brandon Crawford,

Karina Alfaro, Mindi Curran, Clay Markle, and Kellie Eldridge. Michelle Robinson has been with me through it all, since field camp through undergrad thesis then all the trials and celebrations of grad school. Last but definitely not least, I need to thank the people around me every day, who helped with anything and everything, Laurie Marx, Steve

Tillinghast, Colin Wingfield (who made a wonderful adapter to put the tripod on new style benchmarks).

iv TABLE OF CONTENTS

ABSTRACT ...... ii

ACKNOWLEDGEMENTS ...... iv

LIST OF TABLES ...... vii

LIST OF FIGURES ...... viii

INTRODUCTION ...... 1

METHODS ...... 14

Methods for GPS observation of benchmarks ...... 14

Relative sea level change since 1992 based on photo analysis ...... 22

Intertidal organism location comparison methods ...... 27

Urchin pits methods ...... 30

RESULTS ...... 34

GPS observation of benchmarks ...... 34

Relative sea level from photos results ...... 37

Intertidal organism colony location results ...... 38

Sea Urchin location results ...... 39

DISCUSSION ...... 41

GPS observation of benchmarks ...... 41

Discussion of possible Rapid Postseismic Deformation ...... 47

Discussion of relative sea level from photo evidence ...... 53

v Discussion of Intertidal organism location and estimate of relative sea level ...... 54

Urchin pit discussion ...... 59

CONCLUSIONS ...... 61

REFERENCES ...... 62

Appendix...... 67

vi LIST OF TABLES

Table 1. Information about benchmarks included in the survey. Permanent Identification Number (PID), stamp designation, the year each benchmark was monumented, the orthometric height in 1992, measurement error of benchmarks included in Murray et al. (1996) modeling...... 15

Table 2. Results of the GPS survey and comparison to the 1992 leveling to measure interseismic deformation...... 35

Table 3. Results of the intertidal organism position comparison based on photos...... 38

vii

LIST OF FIGURES

Figure 1. Tectonic setting...... 2

Figure 2. Seismicity in the Cascadia subduction zone, Mendocino triple junction, and Juan de Fuca/Gorda plates...... 3

Figure 3. (A) Structure contours showing the upper surface of the subducting Juan de Fuca and Gorda plates (McCrory et al., 2012)...... 4

Figure 4. Location of 1992 M 7.1 main shock and the two M 6.6 aftershocks (From Oppenheimer et al., 1993)...... 6

Figure 5. Uplift along the coast measured using vertical extent of mortality (VEM) of sessile intertidal organisms (from Carver, 1994)...... 7

Figure 6. Vertical extent of mortality (VEM) following the 1992 uplift (from Carver et al., 1994)...... 8

Figure 7. Vertical deformation following the 1964 Prince William Sound earthquake in Alaska and in Japan along the Nankai trough following the 1946 M 8.2 earthquake...... 12

Figure 8.Configuration of the antenna, tripod and receiver I used to complete static GPS observations on each benchmark...... 16

Figure 9.Two types of benchmarks used in this study. Both are first order vertical benchmarks...... 17

Figure 10. Location of benchmarks surveyed and the coseismic deformation model contours from Murray et al. (1996)...... 18

Figure 11. Difference between leveling orthometric height and GPS orthometric height (Modified from Roman, 2007)...... 20

Figure 12. The control point, in red, and the nearby CORS station P160 in yellow...... 21

Figure 13. Elevation change of each benchmark was calculated relative to the Rohnerville control point. The difference between the benchmark’s height in 1992 and 2014/15 is the elevation change...... 22

Figure 14. Relative sea level comparison analysis...... 24

viii Figure 15. Tide level calculation (“Method to find times or heights between high and low waters,” 2013)...... 25

Figure 16. Relative sea level comparison and the effect of sand in mid to upper tidal areas...... 26

Figure 17. Procedure and calculation for using intertidal organism position to measure interseismic deformation via relative sea level change...... 28

Figure 18. Location of the reef south of Mussel Rock where urchin pit measurements were made in the large swash channels. This figure also shows the location of Mussel Rock relative to benchmarks Q229 and Z1465...... 31

Figure 19. Examples of the channel walls with distinct urchin pits...... 32

Figure 20. Schematic showing the method of calculating interseismic subsidence using sea urchin pits...... 33

Figure 21. Interseismic and coseismic elevation change of benchmarks surveyed...... 36

Figure 22. Interseismic elevation change of stable benchmarks with +-10 mm error bars based on GPS survey accuracy (see Discussion section)...... 37

Figure 23. Continuous GPS stations within and near the study area with records shown in IGS08 terrestrial reference frame...... 43

Figure 24. Vertical record of the pair of continuous GPS stations at Cape Mendocino. The yellow lines are horizontal...... 44

Figure 25. Schematic depiction of how the geodetic signal of a minor fault could be affected by an underlying major fault...... 46

Figure 26. Graphical representation of coseismic and interseismic displacement for each benchmark...... 48

Figure 28. Schematic depiction of how rapid postseismic deformation before intertidal organism die-off might change the VEM and uplift estimate...... 53

Figure 29. Examples of how reef morphology could affect the presence of recolonization within a photo extent...... 58

ix Figure 30. Schematic showing how the presence of a wavecut platform between small sea stacks, irregularities or remnants of a higher platform could cause easternmost reef rocks to have smaller colonies, despite being at the appropriate elevation. Although the easternmost rocks might be within the tidal zone, if the waves are very shallow when they reach the rock the water might not support a large colony...... 59

x 1

INTRODUCTION

The Cascadia subduction zone (CSZ), where the Juan de Fuca and Gorda plates

subduct beneath the North American plate, extends from northern California to southern

British Columbia (Figure 1). Convergence rates range from 30 mm/yr in the south to 45

mm/yr in the north (McCrory et al., 2012). The CSZ terminates in the south at the

Mendocino triple junction (MTJ) where the San Andreas fault and Mendocino fault also

terminate (Figure 1). Due to the northward motion of the Pacific plate, along with the

subduction and internal deformation of the Gorda plate, the MTJ is a region of high

deformation rates and seismicity (Figure 1, 2)(Merritts, 1996). Uplifted marine terraces,

both large Pleistocene terraces with 10 – 100 m risers and smaller Holocene terraces with

1 – 3 m risers record deformation of the upper plate in the MTJ region. The geometries of

the subducting Juan de Fuca and Gorda plates have been mapped using intraplate

earthquakes to produce structure contours of the upper surface of the subducting slab

(Figure 3; McCrory et al., 2012). Near the MTJ (Figure 3), the down-going slab flattens beneath an unusually active seismic cluster that is thought to be within the upper plate accretionary wedge (McCrory et al., 2012).

Figure 1. Tectonic setting. The black arrows indicate plate convergence rate and direction. The study area indicated by the yellow box is at the southern end of the CSZ in the region of the MTJ. The Pacific plate has a northward velocity that crushes the Gorda plate and causes it to have an oblique convergence to the North American plate. The exact geometry of the southern CSZ and the MTJ are not well understood. Figure modified from McCrory et al., 2004.

Figure 2. Seismicity (green dots) in the Cascadia subduction zone, Mendocino triple junction, and Juan de Fuca/Gorda plates. The majority of the subduction zone has very little seismicity, however the southernmost end is very seismically active with earthquakes in both the oceanic and upper plate. (Jules Verne Explorer)

Figure 3. (A) Structure contours showing the upper surface of the subducting Juan de Fuca and Gorda plates (McCrory et al., 2012). South of the California-Oregon border the slab is folded, flattening for a few hundred km beneath Eureka and the study area, this may be caused by Gorda’s convergence with the northward moving the Pacific plate. The line C02 – C02’ shows the location and 10 km width of the cross section. (B) The shaded gray area is the Gorda slab as mapped by McCrory et al. 2012. The red star shows the location of the 1992 M 7.1 earthquake. With this interpretation of the slab shape, the 1992 M 7.1 earthquake occurred above the slab interface within the anomalous seismic cluster in the accretionary wedge (McCrory et al., 2012). The approximate location of the M 6.6 aftershocks is shown by the yellow stars (Oppenheimer et al., 1993).

The CSZ megathrust has been historically quiet, with the last large earthquake

being M ~9 in AD 1700 (Atwater et al., 1995; Atwater and Hemphill-Haley, 1997). This earthquake caused a tsunami in Japan that was included in their historical records (Satake et al., 2003). Coastal geologic records indicate that this earthquake caused subsidence along most of the length of the CSZ, as represented by buried marsh soils (Atwater and

Hemphill-Haley, 1997). The geologic record indicates a Holocene history of M 7-9 CSZ earthquakes with recurrence intervals from 200-1200 years (Goldfinger et al., 2012;

Crawford et al., 2014)depending on the interpretation of subduction zone segmentation and earthquake clusters (Graehl et al., 2015).

The only recorded historic earthquake on the CSZ that may have been an interplate megathrust rupture was the April 25, 1992 M 7.1 Petrolia earthquake (Figure

5)(Oppenheimer et al., 1993; Murray et al., 1996). The hypocenter was ~10 km beneath

Petrolia and there was no surface rupture on shore, but it did result in uplift along the coast and produced a tsunami (Oppenheimer et al., 1993; Carver et al., 1994). The magnitude of uplift was estimated using desiccation of intertidal organisms (Figure 5) as well as an National Geodetic Survey (NGS) first order leveling survey (Carver et al.,

1994; Murray et al., 1996). The leveling survey ran along a portion of the uplifted coastline and extended inland to Highway 101. Uplift along the coast was not uniform

(Figure 6); it was greatest at Mussel Rock and tapered to the north and south; no uplift was observed at the Bear River and Punta Gorda. Subsidence was measured inland of the uplift zone (Figure 4). Focal mechanisms and aftershocks showed that the earthquake

occurred on a shallowly east dipping thrust fault approximately 10 km deep

(Oppenheimer et al., 1993). Following the main shock, two M 6.6 aftershocks occurred

offshore within the Gorda plate (Figure 4). Oppenheimer et al. (1993) interpreted these

events to be associated with motion along southeast striking right-lateral strike slip faults

within the Gorda plate.

Figure 4. Location of 1992 M 7.1 main shock and the two M 6.6 aftershocks (From Oppenheimer et al., 1993). The main shock occurred near the town of Petrolia on a shallowly east dipping thrust fault and resulted in the coseismic deformation described by the contours (From Murray et al., 1996). The red solid contours represent uplift in intervals of 0.1 m; blue dashed contours represent subsidence. The aftershocks were left- lateral strike-slip within the Gorda plate (Oppenheimer et al., 1993). Also shown on this map are the geographic locations used in this paper and the major tectonic elements.

Figure 5. Uplift along the coast measured using vertical extent of mortality (VEM) of sessile intertidal organisms (from Carver, 1994). When the patch is lifted out of their habitable zone, the upper organisms die, desiccate and most species turn white. If the uplift is large enough, the whole patch is lifted out of the habitable zone and they all die. In this case, the total patch height is a minimum uplift estimate. If the lower portion of the patch is still alive, then the measurement of uplift based on the height of dead organisms is an accurate estimate of uplift. If the top of the patch extends over the top of the rocks then any measurement based on VEM must be considered a minimum estimate, because the rocks limited the patch’s original height.

Figure 6. Vertical extent of mortality (VEM) following the 1992 uplift (from Carver et al., 1994). The site numbers correspond to the locations shown on the map. The VEM measured is analogous to the amount of uplift at each site. The uplift tapered north and south of the maximum at Mussel Rock. The VEM at each site varied between species.

The source of the 1992 Cape Mendocino earthquake is not well understood. It

occurred on a shallow thrust fault, but it is unclear if it was a megathrust rupture on the interface between the Gorda and North American plates or on a subsidiary fault within the accretionary prism (Oppenheimer et al., 1993; McCrory et al., 2012).The large

9 aftershocks and the large number of smaller aftershocks on the Mendocino fault indicate that there was a release of strain between the Gorda and North American plates during the M 7.1 mainshock (Oppenheimer et al., 1993). This, along with the coseismic crustal deformation, lead to the hypothesis that it may have been a segmented megathrust rupture. If it occurred on the megathrust it indicates that the southern CSZ may be segmented and that rupture on this segmented southern end may or may not initiate an

“unzipping” of the entire CSZ (Personal communication, M. Hemphill-Haley, 2013 -

2015). However, if the 1992 earthquake occurred on a subsidiary fault this does not eliminate the possibility of the CSZ rupture initiating in this highly active southern end and propagating northward.

Although the CSZ has been historically quiet, other subduction zones around the world have had moderate to large earthquakes that have been recorded with modern instrumentation and can give us insights to how the CSZ might behave during its next earthquake. To determine if the 1992 Cape Mendocino earthquake was a segmented megathrust rupture, I am comparing my measurement of postseismic or interseismic deformation to that which has occurred on other subduction zones following historic earthquakes. Because a subduction zone is the boundary between two plates, there should be post and inter seismic deformation as the stress builds between the colliding plates

(Wang et al., 2012). Although there have been many historic subduction zone earthquakes, the vertical postseismic and first few decades of interseismic deformation has only been studied on a few. The Alaska 1964 M 9.2 and a pair of M 8+ earthquakes in Nankai, Japan in the 1940’s have the most extensive vertical deformation records

10 during the first few decades following the earthquakes (Cohen and Freymueller, 2004;

Savage and Thatcher, 1992). Although the many M 7-8 earthquakes on other subduction zones might be more analogous to the 1992 Petrolia earthquake, the vertical interseismic deformation is not well studied beyond the postseismic transient. Perhaps this is due to the smaller magnitude of deformation following a smaller earthquake making measurement more difficult and a lack of precise instrumentation in these remote regions.

Postseismic and interseismic deformation observed in Alaska and Japan can provide a frame of reference to interpret the interseismic deformation measured in this study. In both Alaska and Nankai, deformation during the first few decades occurred at a very fast rate (Figure 7). The postseismic deformation was in the opposite direction as the coseismic deformation. In Alaska the deformation rate during the first year following the

1964 earthquake was 80-90 mm/yr, and 30 mm/yr averaged over the 11 years following the earthquake, resulting in 0.37 m of uplift in only 11 years where the maximum coseismic subsidence had been about 2 m (Cohen and Freymueller, 2004). During these authors’ 11 year study, the postseismic deformation may not yet have been completed, as an interseismic deformation rate of 30 mm/yr is much greater than the 4.8 mm/yr uplift rate observed prior to 1964 (Cohen and Freymueller, 2004). It is important to note that the maximum measured coseismic uplift in the 1964 earthquake was ~10 m, but the interseismic deformation of the uplifted regions is not well studied. In Nankai the uplift record is much longer and postseismic deformation was completed in about 10 years, then the deformation rate became more constant at about 5 mm/yr (Savage and Thatcher,

1992). Both of these other subduction zones had very high deformation rates in the first

11 couple decades resulting in a large portion of the coseismic deformation being recovered.

This fast postseismic deformation has been observed on many subduction zones, and the time span varies from days to years before the steady interseismic deformation rate is reached (Melbourne, 2002). The cause of the postseismic deformation is often attributed to viscoelastic relaxation of the mantle following the earthquake (Ueda, 2003; Wang et al., 2012), whereas the slower interseismic deformation rate is attributed to strain accumulation in the upper plate due to plate convergence (Wang et al., 2012).

Figure 7. Vertical deformation following the 1964 Prince William Sound earthquake in Alaska and in Japan along the Nankai trough following the 1946 M 8.2 earthquake. In both systems the deformation rate is fastest during the first year after the earthquake and quickly decays to a slower rate that is close to the interseismic deformation rate observed before each historic earthquake. A) Cumulative vertical deformation in Alaska in 1965, 1968, and 1975 (Brown et al., 1977). B) Profile showing location and magnitude of coseismic and postseismic deformation in Alaska relative to distance from the trench (Cohen and Freymueller, 2004). C) Uplift rate in Alaska vs. years following the earthquake, showing the decay of deformation rate during the decade following the earthquake. Data used in this graph is from Brown et al. (1977). The red line shows the rate during the years covered by each leveling survey. Blue line is the power function fit line that shows the fast decay of deformation rate over the first 2-4 years. D) Profiles of vertical deformation rate perpendicular to the Nankai trench (Thatcher, 1984). In the first year after the 1947 earthquake 1/5th of the coseismic deformation had been recovered (Thatcher, 1984).

If the 1992 earthquake occurred on the megathrust, I expect to observe cumulative deformation on the scale of at least 10 – 20 cm. This is proportional to the observed deformation in Alaska, where 2 m of coseismic subsidence had 0.37 m of postseismic uplift in 11 years (Zweck et al., 2002). I estimated this under the assumption that the relationship between coseismic and postseismic deformation, with postseismic deformation in ~10 years being 18% of the coseismic deformation, is the same in any subduction zone earthquake regardless of earthquake magnitude. However, if the earthquake occurred on a subsidiary fault I expect to observe much less than 10 cm of deformation, as there would be much less viscoelastic relaxation in a crustal fault and the fault should have a much lower strain rate than the megathrust.

METHODS

To measure deformation in the 23 years since the 1992 earthquake, I used two different data sets and three different methods of analysis. First, I used GPS benchmark observations and a high precision experimental geoid to compare the current elevation of select benchmarks to their 1992 elevation in order to measure vertical deformation.

Second, I had access to a set of photographs taken during the intertidal organism die-off

study of 1992 following the co-seismic uplift event. I used these photos in two ways, first

to compare relative sea level directly by identifying and comparing distinctive rock

features, and second to compare positions of intertidal organism colonies as a proxy for

relative sea level change.

Methods for GPS observation of benchmarks

Following the 1992 earthquake sequence near Petrolia, uplift was identified by intertidal organism die-off (Carver et al., 1994), a NGS first-order leveling survey was conducted to measure coseismic deformation (Stein et al., 1992; Murray et al., 1996).

The 1992 leveling was compared to surveys conducted in 1935 and/or 1942 where data

was available (Stein et al., 1992; Murray et al., 1996). During the 1992 survey, new

benchmarks were emplaced and included in the leveling (Table 1). Using benchmarks

selected based on suitability for satellite observation and distribution across the coseismic

15 deformation area, I compared current benchmark orthometric height to 1992 orthometric height to measure crustal deformation over 23 years.

1992 1992 Year orthometric leveling PID Stamp Monumented height (m) error (mm) LV0661 N1401 1992 76.301 LV1251 Z1465 1992 9.261 LV0368 P229 1935 9.169 3.53 LV1253 RMN01 1936 10.676 LV0366 M229 1935 93.88 LV1263 M1467 1992 157.38 LV0405 R649 1942 681.615 5.67 LV0404 G275 1935 690.79 2.63 LV0410 K275 1935 614.051 6.62

To achieve precise benchmark positions, I used a Trimble Zephyr Geodetic antenna with a Net R9 receiver provided by UNAVCO on a heavy duty GPS tripod with a center leg (Figure 8). Individual site observations were 6 hours long; each benchmark was observed with this instrumentation one time. The antenna was oriented to true north and the same equipment was used for all observations. Because of time constraints I was unable to observe the benchmarks multiple times for redundancy.

Figure 8.Configuration of the antenna, tripod and receiver I used to complete static GPS observations on each benchmark. The antenna was mounted on the GPS tripod belonging to HSU Geology that is usually used for the RTK base station. This mount consists of three stabilizing legs and a fourth center leg that is directly beneath the center of the antenna. This leg is placed on the benchmark and leveled using an attached bubble level.

Benchmarks suitable for satellite observation were located using NGS data sheets.

The current satellite observation suitability was assessed based on the size of obstructions greater than 10 degrees above the horizon. Slope stability was assessed for each benchmark and if the benchmark was disturbed or within a landslide it was not included in the survey. The benchmarks were all first order vertical benchmarks and were of two physical types depending on age; I did not discriminate between benchmark types and

17 observed a few of each (Figure 9). I selected 8 benchmarks for observation that had few obstructions and were well distributed across the coseismic deformation zone (Figure 10).

Figure 9.Two types of benchmarks used in this study. Both are first order vertical benchmarks. (A) “old” type, metal disk set in a concrete post. (B) The pointed tip of the tripod center leg fits into the benchmark divot. (C) “New” type of benchmark, consisting of a metal rod protected by a sleeve and cap. The actual benchmark point is the cone- shaped tip of the metal rod. The cap is set flush with the ground surface and the rod is usually a few inches beneath the ground surface (Smith, 2010). (D) Because the benchmark is pointed, the pointed tip of the tripod cannot be stably placed on it. I used a custom-made tripod tip that fits over and around the benchmark rod. The custom tip has the same effective height as the default tip, so no height adjustment factor needs to be used.

Figure 10. Location of benchmarks surveyed and the coseismic deformation model contours from Murray et al. (1996). The chosen benchmarks covered both uplift and subsidence zones, with the farthest south, Z1465, being proximal to Mussel Rock, where the highest coseismic uplift was measured using intertidal organism die-off. N1401 is the control point in Rohnerville, to which all my measurements were compared to calculate interseismic deformation. The control point is well outside the 1992 deformation zone as reported by Stein et al. and Murray et al. (1992; 1996).

Benchmark survey data in the form of RINEX files were then submitted to

NOAA’s Online Positioning User Service (OPUS), which averages the coordinates

submitted between three single-baseline solutions from nearby continuous GPS (CORS)

stations (OPUS). I then input the corrected ellipsoid coordinates from OPUS into

xGEOID14B to get an orthometric height. Because leveling and GPS measure height in

fundamentally different ways (Figure 11), and the geoid model is the theoretical, modeled

surface that rectifies the two (Figure 11b). I chose xGEOID14B because it incorporates

thorough aerogravity measurements conducted in 2011 and covering the entire study area

(National Geodetic Survey). This geoid is being created to define a new vertical datum and is not referenced to a previous datum such as NAVD88 (Roman, 2014). Because the

1992 leveling orthometric heights are in NAVD88, I use a control point outside the deformation area to compute benchmark elevation changes from 1992 to 2015 relative to the vertically stable control point. I chose benchmark N1401 in Rohnerville as my control point, because it is outside the modeled deformation area (Murray et al., 1996). Nearby continuous GPS station, P160, shows no vertical motion (Figures 10, 12) for the 10 years it has been active. Thus, it appears that the station has remained vertically stable for the past 23 years.

Figure 11. Difference between leveling orthometric height and GPS orthometric height (Modified from Roman, 2007). A) GPS measures height relative to the ellipsoid, a geocentric mathematical approximation of the idealized shape of the earth. Leveling measures height relative to the geoid, a natural gravitational equipotential surface that undulates due to geologic factors such as variations in density and thickness of the earths crust (Roman, 2007). B) To rectify the two measurement methods, an orthometric correction is applied to GPS measurements using a model of the geoid. The closer the geoid model is to the real undulations of the equipotential surface, the more closely GPS orthometric height will replicate orthometric height determined through leveling. The accuracy of the geoid model is imperative for accurate comparison of leveling derived and GPS derived elevations (Cohen et al., 1995).

Figure 12. The control point, in red, and the nearby CORS station P160 in yellow. This map shows the proximity of the control point to the continuous GPS, they are approximately 1 mile apart. The orange lines are faults from the USGS (USGS), and they show that there are no mapped faults between the control point and the P160, thus they should be experiencing the same tectonically induced motion. The control point is also on flat ground away from the edge of a large terrace, so it is unlikely to be experiencing slope failure. The plot shows the vertical position change in mm of station P160 for the 10 years since its installation. The position has fluctuated around zero, likely due to seasonal factors (Dong et al., 2002; Pan et al., 2015) but there has been no overall change in P160’s elevation in IGS08, a terrestrial reference frame.

To calculate the elevation of each benchmark relative to the control point, N1404 in Rohnerville, I differenced each benchmark’s 1992 leveling-derived orthometric elevation (Table 1) with the 1992 height of the control benchmark, N1401 in Rohnerville

(Figure 13). I then repeated this procedure with the 2015 benchmark height and

Rohnerville benchmark height, both derived from my GPS observations and processed through OPUS and GEOID14b. The result provides a 1992 and a present orthometric

height of each benchmark relative to the Rohnerville benchmark. Comparison of these

relative heights yields benchmark elevation change since 1992, similar to the results if

there had been a 2015 leveling survey.

Figure 13. Elevation change of each benchmark was calculated relative to the Rohnerville control point. The difference between the benchmark’s height in 1992 and 2014/15 is the elevation change.

Relative sea level change since 1992 based on photo analysis

The second data set I used to measure elevation change incorporated photos from

Dr. Angela Jayko and Dr. Bob Rasmussen taken during the Vertical Extent of Mortality

(VEM) survey of Carver et al. (1994).

I compared the relative water level along coastal exposures between 1992 and present using intertidal rocks as relative elevation markers. I digitized the photos from

1992, post earthquake photographic slides and analyzed them using Adobe Photoshop, then used them for both the intertidal colony location and relative sea level studies. To accomplish this it was necessary to capture the same photo angle and tidal level. First, I identified the location from which each picture was taken by identifying unique rocks and

landmarks in the field. I recorded this position as a GPS coordinate for later reference.

Then, I found the date that the picture was taken from Dr. Jayko’s field notes based upon general location because they visited a different location each day (Mussel Rock, Devils

Gate, Sea Lion Rock, etc.). Using Dr. Jayko’s field notes, observation of whether the picture was taken in the afternoon or morning, and historic tidal records from Shelter

Cove (CeNCOOS, 2014), I was able to determine the approximate time that each photo was taken. Then I used a solar calculator (Colleti, 2011) to estimate the sun azimuth and elevation at that time (Figure 14). I then used a high resolution coastal LiDAR DEM to model the shadows in ESRI ArcMap using the hillshade tool with a model shadows option selected. I then compared the modeled shadows to shadows in the pictures. If shadows didn’t match I recalculated the sun position and remodeled the shadows at a later or earlier time of day as needed. Using this method I estimated a time of day for when the photo was taken. If the photo was not taken at low tide, I used a tide calculation chart (“Method to find times or heights between high and low waters,” 2013) to calculate the tide level during the photo (Figure 15). Finally, I attempted to capture a present day version of the photo on a day with the same tidal level. With these two pictures, I could compare the relative sea level based on water level on the intertidal rocks.

Figure 14. Relative sea level comparison analysis. A) The original 1992 photo, #125 from Dr. Angela Jayko. Photographer position, shown by the green dot in B, was found via field analysis and comparison of visible distant hills. B) The viewshed from the photographer position confirms correct portions of distant hills are visible and allows for comparison of shadow locations on those hills. C) The sun angle at 8:30 am results in shadows that cover the photographer and most of the beach, clearly not agreeing with the original photograph. D) The shadows resulting from the 9:00 am sun angle reach just to the base of the hills and cover appropriate areas of the distant hills. The shadows at 9:00 agree with the original photograph, giving a photograph time to within approximately 30 minutes.

Figure 15. Tide level calculation (“Method to find times or heights between high and low waters,” 2013). If the photo was not taken at low tide, I used this chart to calculate the level of the tide at the time the photo was taken. The height and time of the previous and following tide are required. I could then find the height of the tide at any time. This chart can also be used to find the time of any particular tide, which is useful in recreating the 1992 photos.

The difficulties of this methodology were threefold. The photo angle had to be recreated nearly perfectly. Where the current sand level was higher than the 1992 sand level it sometimes made it impossible to get the appropriate photo angle. In some cases, high sand also obscured rock features or filled in areas that were open to the waves in

1992 (Figure 16). Finally, the size and direction of waves and wind could affect apparent relative sea level on the decimeter scale; historic records are insufficient to quantify these factors.

Figure 16. Relative sea level comparison and the effect of sand in mid to upper tidal areas. A) this set of photos was taken south of the mouth of the Mattole River. Using tide records from Shelter Cove (CITE) and the tide level calculation in figure 17, I determined that these photos were taken at the same tide level. Although the angle is not perfect, you can see that the water level relative to the rocks in the 2015 photo is within 10 cm of the water level in 1992, indicating little to no relative sea level change. B) This photo set was taken near Devils Gate. In 1992 the sand level was much lower than in 2015, allowing ocean water to reach farther up the beach. These photos were not taken at the same tide level. In this instance and a few others, high sand levels in 2015 prevented comparison of relative sea level.

Intertidal organism location comparison methods

After the 1992 earthquake, coseismic uplift was measured using the changed vertical levels of desiccated intertidal organisms (Carver et al., 1994). Because intertidal organisms live on the rocks within a particular tidal range the vertical distribution of a species can be used as an indicator of relative sea level; any type of sessile intertidal organism can be used. Within a few days of the 1992 uplift, intertidal species no longer within their former habitable tidal zones (Carver et al., 1994). By measuring the height of the desiccated portion of a species, the coseismic uplift was estimated. Post-seismic or interseismic deformation can be measured using relative sea level change since the coseismic uplift. Slow interseismic relative sea level change can be estimated by comparing current species locations to their pre-coseismic uplift locations (see Figure 17 for methodology).

Figure 17. Procedure and calculation for using intertidal organism position to measure interseismic deformation via relative sea level change. The top of an intertidal species distribution is controlled by desiccation and is directly and consistently related to sea level. When the rocks uplift out of the water, the upper portion of the population dies and a new upper limit is established. By measuring the distance on the rocks between the pre- uplift distribution top and the top of the reestablished species distribution (D), and subtracting this value from the measured coseismic uplift (U), the interseismic deformation is measured (S). If S is negative, there has been interseismic subsidence, if it is positive there has been interseismic uplift.

To measure the difference between post-1992 uplift event species distribution and present position I used 1992 field photos acquired by Dr. Angela Jayko and Dr. Robert

Rasmussen. I then took comparable photographs of the same locations for measurement.

First, I scanned and examined the 1992 photos to identify rocks with prominent species that had sharp distribution boundaries. I then enhanced the most promising photos in

Adobe Photoshop by manipulating the contrast, brightness, and color in order to better identify the species and to delineate the edges of the population as precisely as possible.

Photo locations for population comparison were selected that had at least one identified

species with a well-defined upper distribution. Once I completed preliminary

examination of the 1992 photos I went to the field area during very low tides to recreate

the photos. While in the field, I tried to recreate the photo angle as accurately as possible.

Back in the lab, I marked the population edges on current photographs using Adobe

Illustrator. I made the measurements within the photos using the in-photo scale because it

simplified the identification of rock features and allowed for faster data collection during

very low tides by eliminating the need for detailed surveying of each rock face.

I compared photographs of the 1992 species locations to the 2014/15 locations. I

targeted the upper limit of the species because it is controlled by desiccation. The upper

limit can be artificially high in high-splash zones where the spray allows organisms to live higher on the rocks (Melnick et al., 2012). The lower limit is often controlled by predation and, while it may represent the desiccation point of the predating species, there may be other complicating factors (Melnick et al., 2012; personal communication,

Rasmussen, 2015). Using the in-photo scale, I then measured the difference between the top of the present population and the desiccated 1992 top which represented the pre-uplift

habitable zone. Importantly, if the population top in either photo extended over the top of

the rock the true upper extent may not have been reached and the measured distance (D)

between colony upper limits is a minimum, and coseismic uplift or interseismic

subsidence could be underestimated.

Urchin pits methods

Sea urchins are mobile intertidal organisms that move to higher or lower positions on intertidal rocks as sea level changes (Figure 20)(Rasmussen, 2105). Because they are mobile, they were not used in coseismic uplift measurements by desiccation. However, when long standing colonies relocate they leave behind pits in the rock, which can be used to identify sea level change (Personal communication, R. Rasmussen, 2015). The pits are half-circle shaped indentations in the rock that are approximately 5 – 10 cm in diameter. After the 1992 uplift event, sea urchins in certain swash channels south of

Mussel Rock (Figure 18) recolonized lower on the rocks and formed a distinct zone of fresh new pits (Figure 19) (Personal communication, R. Rasmussen, 2015). Above the fresh pits there is a zone approximately 20 cm high, of older weathered pits that were already vacant before the 1992 uplift (Personal communication, R. Rasmussen, 2015). I estimated the relative sea level change since 1992 by measuring the distance between the top of the current urchin colonies and the top of the pits vacated after the uplift (Figure

20). If the distance between the urchins and top of the old pits was less than the uplift measured in that location then there has been relative sea level rise associated with crustal subsidence since 1992.

Figure 19. Examples of the channel walls with distinct urchin pits. Photos are not from the same location, although they are on the same reef immediately south of Mussel Rock (Figure 19). The “old” pits, occurring within a zone of ~20 cm, were not occupied immediately before the 1992 uplift (Personal communication, R. Rasmussen, 2015). Empty pits were occupied before the uplift and vacated in the months following the earthquake (Personal communication, R. Rasmussen, 2015). Occupied pits may have been occupied before the uplift and are occupied today. They indicate the new tide level suitable for the sea urchins. Empty pits observed in 2015 are covered with various upper tidal algae.

Figure 20. Schematic showing the method of calculating interseismic subsidence using sea urchin pits. A) Pre-1992 uplift condition, a zone of older empty pits with occupied pits below them. B) Conditions after urchins have recovered from the uplift and recolonized with respect to the new, lower sea level. The height of the zone of empty fresh pits is equal to the uplift. C) If subsidence has occurred, sea level would rise and the zone of unoccupied pits will be less than the uplift measured in 1992. D) If there has been no subsidence since 1992 the height of empty pits will be equal to the uplift, and will also equal the height of empty pits after the colonies have reestablished following the uplift.

RESULTS

This study resulted in various estimates of vertical benchmark and landmark elevation change relative to two different datums. In the first instance, GPS measurements resulted in elevation change estimates relative to the control point located outside the coseismic deformation area. Relative sea level, estimated from photo time series, intertidal organism colony locations, and sea urchin pit analysis provided estimates of elevation change relative to average local sea level.

GPS observation of benchmarks

GPS derived elevation changes for this study are relative to the control point,

benchmark N1401 in Rohnerville (Figure 10), located outside the coseismic deformation

area of the 1992 earthquake. According to nearby continuous GPS station P160

benchmark N1401 has been vertically stable for over 10 years. When the measured GPS

sites are normalized to the Rohnerville control N1401 the result is that there has been less

than 30 mm of elevation change since the leveling in 1992 (Table 2). In general, benchmarks within the coseismic uplift zone have subsided and the benchmarks within the coseismic subsidence zone have uplifted (Figure 21). Benchmarks that are located where the greatest coseismic deformation occurred correspond with the greatest interseismic deformation in the opposite direction. Benchmarks near the modeled coseismic uplift-subsidence hinge (P229, RMNO1, M229 in Figure 10) have variable

35 vertical interseismic deformations near zero (Murray et al., 1996). The signal is small and near the level of measurement accuracy (ca. 1-2 cm), but it is detectable (Figure 22).

Table 2. Results of the GPS survey and comparison to the 1992 leveling to measure interseismic deformation.

Elevation xGEOID14B change 2014-15 1992 relative to Elevation orthometric orthometric Rohnerville change rate PID Stamp height (m) height (m) (cm) (mm/yr) Control LV0661 N1401 75.354 76.301 Point LV1251 Z1465 8.335 9.261 -2.1 -0.91 LV0368 P229 8.222 9.169 0 0.00 LV1253 RMN01 9.732 10.676 -0.3 -0.13 LV0366 M229 92.927 93.88 0.6 0.26 LV1263 M1467 156.573 157.38 -14 -6.09 LV0405 R649 680.655 681.615 1.3 0.57 LV0404 G275 689.849 690.79 -0.6 -0.26 LV0410 K275 613.078 614.051 2.6 1.13

Figure 21. Interseismic and coseismic elevation change of benchmarks surveyed. Benchmark M1467 is not included in this figure because it is unstable. The benchmarks are arranged based on modeled or measured coseismic elevation change, with the greatest uplift on the left and most subsidence on the right. Blue diamonds show the interseismic elevation change measured relative to the control point, benchmark N1401 in Rohnerville. Orange triangles are coseismic elevation change as measured during the 1992 leveling survey (Stein et al., 1992; Murray et al., 1996). The red square shows the elevation change at Mussel Rock as measured through vertical extent of intertidal organism mortality (Carver et al., 1994). This coastal location is proximal to benchmark Z1465 (Figure 19) and may represent coseismic change at that benchmark had it been

37 installed prior to 1992. This comparison shows the large difference in magnitude of coseismic and interseismic deformation.

Figure 22. Interseismic elevation change of stable benchmarks with ±10 mm error bars based on GPS survey accuracy (see Discussion section). The benchmarks are arranged from greatest coseismic uplift to greatest coseismic subsidence along an arbitrary axis. Z1465 is close to Mussel rock, where the greatest coseismic uplift was measured (Figure 12). Z1465 has experienced the most subsidence since 1992. K275 experienced the greatest coseismic subsidence, and has also had the greatest interseismic uplift. Benchmark M1467 has not been included on this graph due to instability (see Discussion section).

Relative sea level from photos results

Although I identified the location and time on many photos, I was only able to complete relative sea level measurements south of the mouth of the Mattole river and at

Cape Mendocino. In most locations the tide level I estimated was lower than any tides that would occur during daylight over the course of my study, so they could not be

38 replicated. At those two locations my estimate of relative sea level change was within

~10 cm of 1992 postseismic levels.

Intertidal organism colony location results

At Mussel Rock I made comparisons of at least one intertidal species in each of 5 photo locations (Table 3, Appendix A). Despite all being within a 0.1 km2 area on the reef at Mussel Rock, the results of these measurements are highly variable, ranging from

+0.3 to -0.6 m. The average absolute land level elevation change is -0.26 m.

Table 3. Results of the intertidal organism vertical position comparison based on photos. Photo name corresponds to whether it is from the Rasmussen (BR) or Jayko (AJ) collection (Appendix B). Coseismic uplift estimate at Mussel Rock from Carver et al. (1994). The distance between tops is measured by the top of the population in 1992 vs. 2014/15. It represents relative sea level change from before the earthquake to present. The absolute elevation change is the interseismic elevation change of the rocks relative to sea level calculated by subtracting the distance between tops from the coseismic uplift. “+” represents interseismic uplift while “–“ represents interseismic subsidence. Absolute distance coseismic elevation Photo between tops uplift (cm) change since (cm) 1992 (cm) BR 6 140 120 -20 BR 8 140 170 30 BR 15 140 180 -60 AJ 73 140 100 -40 BR 12, 13 140 100 -40 and 16

I used mussels and barnacles for elevation comparison in most locations because

it was difficult to differentiate among the seaweed species in the 1992 photos. Mussels and barnacles formed thick, well defined bands that were easily identified and formed sharp boundaries. However, in most of the photo locations the current mussel population is scarce and does not form the thick, well defined bands. In these cases, I made measurements using what mussels were present, though they were mostly living in cracks in the rocks or on faces with high splash.

Sea Urchin location results

I made sea urchin measurements in three swash channels in the reef south of

Mussel Rock (Figure 18). Results are highly variable, however the average is nearly equal to the coseismic uplift, suggesting, based on these measurements, that there has been little to no land level change (Table 4).

Table 4. Measurements of distance between to top of the current urchins and fresh urchin pits. These measurements were conducted at the reef south of Mussel Rock in unobstructed swash channels with vertical walls and identifiable empty pits. Top of fresh pits to Location top of urchins (cm) 1 110 1b 100 1c 110 1d 90 1e 110 2 150 2a 130 2b 200 2c 190 3 180 3a 160 3b 170 Average 142

DISCUSSION

GPS observation of benchmarks

The GPS methods as applied here allow for 10 mm of precision in the 2015 benchmark vertical positions (Personal communication, B. Hammond, 2015). Potential sources of error within the GPS measurements include mechanical imprecision (less than

1 mm in the geodetic antenna), multipath interference, geoid errors. Errors in the elevation change calculation could also be introduced through the 1992 leveling; leveling errors were published in Murray et al. (1996) for some of the benchmarks included in my study (Table 1). The error for leveling ranges from 3 mm up to 7 mm for individual benchmarks. Rounding the 1992 benchmark position error to 10 mm and the GPS vertical error to 10 mm, I use a vertical elevation change error of 20 mm.

One surveyed benchmark had anomalously large subsidence, suggesting that it is unstable. Benchmark M1467 (Figure 10) had 140 mm of subsidence (Table 2), over four times the apparent subsidence of any other benchmark in the study. This benchmark is located along the Mattole road near the mouth of the Bear River on a steep hillside with many slumps and small landslides. Although the benchmark itself appeared stable (the rod was centered within the casing and the cap was not covered with excess soil), the anomalously large amount of subsidence and the unstable landscape surrounding the site leads me to conclude that it is likely moving because of slope processes.

A sparse network of continuous GPS stations is within the study area (Figure 23).

The vertical velocity of these stations does not agree with the benchmark displacements that I measured at nearby benchmarks. The largest difference is between benchmark

M229 and the GPS stations at Cape Mendocino. There are two stations at Cape

Mendocino, CME5 and CME6, standard installation procedure for Coast Guard reference stations (National Geodetic Survey). The continuous GPS stations and M229 are less than

1 km apart and on the same hill, yet my measurement of M229 indicates 6 mm of uplift since 1992 (essentially zero, considering ±10mm error), and since installation in 1996, the Cape Mendocino GPS stations indicate an average subsidence rate of ~3 mm/yr.

However, they appear to only subside during earthquakes, and the velocity between earthquakes is ~0 (Figure 24). The regular oscillations in the vertical record are likely due to seasonal factors (Dong et al., 2002). Many of the earthquakes associated with station subsidence should not have affected the area tectonically as they are on distant faults or within the Gorda plate. The stations are located on top of the ocean bluff at Cape

Mendocino, with highly deformed Franciscan Complex bedrock and geomorphic evidence of active mass wasting with wave erosion at the base. The bluff may be unstable and subsiding due to slope processes or gravitational collapse during seismic shaking.

Figure 23. Continuous GPS stations within and near the study area with records shown in IGS08 terrestrial reference frame.

Figure 24. Vertical record of the pair of continuous GPS stations at Cape Mendocino. The yellow lines are horizontal. Between earthquakes, indicated by the dashed gray vertical lines, the station velocity averages about 0. During earthquakes the stations subside, even when the earthquake is far away and should not cause tectonic motion at Cape Mendocino.

A similar pattern of subsidence is observed at the Bear River Ridge station

(P159), while the Punta Gorda station (P157) appears to be very stable (Figure 23). The

Monument Ridge station (P158), which is the most proximal to benchmark K275,

45 appears to be uplifting at up to 4 mm/yr, however this was due to the growth of a tree close to the antenna creating a multipath error (Puskas, 2015).

The rate of vertical deformation differed for each benchmark. Z1465, near Mussel

Rock, subsided at ~1 mm/yr (Table 2). Benchmark K275 has uplifted ~1 mm/yr (Table

2). This is low for the current interseismic deformation on the Cascadia margin, which ranges from 1 – 4 mm/yr of uplift (Burgette et al., 2009). In comparison to postseismic motion on other subduction zones with historic earthquakes, this rate is also small. For example, in Alaska the first 11 years after the 1964 M 9.2 earthquake experienced 5 cm/year of uplift (Zweck et al., 2002) and Nankai Japan there was ~10 mm/year of uplift for the first 16 years following the 1946 M 8.2 earthquake(Savage and Thatcher, 1992).

The measured vertical interseismic deformation is does not fully mirror the coseismic deformation. As an example, benchmark P229 had 0.7 m of coseismic displacement, yet I can detect no interseismic elevation change. If the 1992 earthquake did not occur on the subduction zone, but the subduction zone is still active beneath the study area, deformation from both systems may be expressed in geodetic measurements

(Figure 25). It is possible that, if the 1992 earthquake occurred within the North

American plate (NAP) interseismic deformation from the 1992 earthquake could cause benchmark P229 to be subsiding, yet at the same time deformation from the underlying subduction zone could be uplifting part or all of the study area. If 1992 subsidence and

CSZ uplift have the same rate, they would cancel and there would be zero net change in elevation at benchmark P229. However, at benchmark Z1465, at location of maximum, on-land, coseismic deformation, interseismic deformation associated with the 1992

earthquake would be expected to be greater and may be outpacing CSZ uplift, causing net

subsidence at Z1465.

Figure 25. Schematic depiction of how the geodetic signal of a minor fault could be affected by an underlying major fault. The two fault systems could have conflicting deformation fields, causing the geodetic signal to show the net elevation change that fails to fully represent either system.

If the 1992 earthquake occurred on the interface between an abandoned slab

within the accretionary prism and the North American plate (NAP), as proposed by

McCrory et al. (2012), then it would still allow for the stress change in the Gorda plate that initiated the two aftershocks. Presumably, if the abandoned slab is between the upper

plate and the Gorda plate, then there must be coupling between the abandoned slab and

the Gorda plate. The movement of the 1992 event on the slab/NAP interface would have

allowed movement of the Gorda plate relative to the NAP, even though the interface

between the Gorda plate and the abandoned slab, which would be the active subduction

zone interface, did not rupture.

Discussion of possible rapid postseismic deformation

Within the coseismic subsidence region, benchmark K275 (Figures 10, 21, 22)

had the greatest interseismic uplift. This does not agree with the deformation modeled by

Murray et al. (1996), which shows the greatest subsidence located farther southwest

(Figure 10). Murray et al. (1996) suggests that errors in their model may be due to

calculation parameters that include a simple rectangular rupture area and evenly

distributed slip along the entire rupture area. Because leveling conducted in 1935 or 1942

occurred several decades prior to the 1992 earthquake and subsequent re-leveling, the coseismic deformation measurements might include some interseismic deformation from the previous earthquake cycle.

Figure 26. Graphical representation of coseismic and interseismic displacement for each benchmark. Some benchmarks were placed in 1992, thus do not have a coseismic displacement. Benchmark P229 had 0 mm of interseismic displacement. The coseismic and interseismic displacements are shown at different scales. This is done for graphical reasons due to the order of magnitude difference in displacements.

In comparing the coseismic uplift estimates from intertidal organism desiccation and benchmark leveling, one must consider error in estimates. Benchmark P229 (Figure

10), which is along the coastline within the Devils Gate reef, where coseismic uplift was measured using vertical extent of mortality (VEM) of intertidal organisms. The VEM

uplift estimate was 1.1 m with standard deviation (std) = 0.2 m. Benchmark P229 had an

uplift measurement of 0.7 m. The 0.4 m discrepancy is just within 2 std of the VEM measurement, so is possibly explained by measurement errors. This discrepancy also exists between Mussel Rock and the nearby benchmark Q229 (Figure 18). The VEM derived uplift estimate at Mussel Rock was 1.4 m with a standard deviation of 0.3 m

(Carver et al., 1994). Benchmark Q229, ~2 km north of Mussel Rock, had a coseismic

uplift of 0.88 m estimated from the leveling (Murray et al., 1996). Again, the benchmark

uplift is just within the 2 std measurement error of the nearby VEM uplift and possibly

explained by measurement errors. However, I also consider the implications of whether

the measurement discrepancy is actually real (Figure 27). Murray et al.’s (1996) model A

included leveling west of Highway 101. Murray et al. (1996) also created deformation

models using only the VEM coastal displacements and horizontal GPS displacements.

Model A had much larger normalized root mean square (nrms) error than the models that

did not include the leveling. Murray et al., (1996) propose that this difference in nrms

could indicate that the leveling coseismic displacements may include some deformation

from the previous earthquake cycle. Carver et al. (1994) completed the VEM

measurements within 1 month of the earthquake (Carver et al., 1994). Any pre-seismic or

previous cycle interseismic deformation would have been accommodated by organisms

slowly recolonizing, potentially making the VEM measurements a very pure coseismic

signal. The GPS measurements conducted in 1989, 1991, and in May 1992, could

represent a pure coseismic signal with little interseismic or postseismic deformation. The

models that exclude the leveling west of Highway 101 have lower nrms (Murray et al.,

1996). The leveling along the coast at P229 was done in either 1936 or 1942 (resurveying of an unspecified portion was conducted in 1942 to correct a systematic error in the 1936 survey) and in October 1992, 6 months after the earthquake (Murray et al., 1996). The difference in benchmark elevation between the 1936/1942 and 1992 surveys was considered to be the coseismic displacement, for P229 it was 0.7 m. With the large interval between the 1936/1942 and 1992 leveling, the coseismic signal could be contaminated by an interseismic signal from the last 50 years of the previous earthquake cycle (Murray et al., 1996). This could be a cycle of repeated 1992-like earthquakes or the megathrust rupture of the CSZ in 1700. If an interseismic signal is the cause, the 0.4 m discrepancy would be explained by ~0.8 cm/year of subsidence prior to the 1992 earthquake. This rate is an order of magnitude larger than the interseismic deformation I have observed even at the extremes of my data. If previous cycle interseismic deformation is used as an explanation for the discrepancy, then the interseismic behavior in the first 20+ years of this earthquake cycle differs greatly from the end of the previous earthquake cycle.

Figure 27. Schematic representation of scenarios in which the coseismic intertidal organism VEM uplift measurement is 0.4 m greater than the leveling derived coseismic uplift measurement at benchmark P229. The discrepancy may be explained by interseismic deformation prior to the 1992 earthquake (dashed line). It might also be due to rapid postseismic subsidence in the 6 months between the earthquake and the leveling survey. An example of possible exponential decay of postseismic deformation is shown in the expanded window.

Another possible explanation for this 0.4 m discrepancy could be rapid postseismic subsidence in the 6 months between the earthquake and the post-event leveling. If postseismic deformation occurred linearly for 6 months there would have been ~6.6 cm/month of subsidence. This would have had to stop immediately after the

1992 leveling, because there has been a maximum of 1 cm deformation in the entire 23 years since that leveling. However, if postseismic deformation was rapid immediately after the earthquake and decayed exponentially over the next 6 months, it could have been nearly 0 at the time of the 1992 leveling, and continued to be 0 for the next 20+ years. Extremely rapid postseismic movements that decay to the interseismic rate have been observed on other subduction zones, with postseismic transients being measured for days to months to years after an earthquake (Melbourne, 2002). However, rapid postseismic deformation could potentially affect the VEM measurements. Any post seismic subsidence that occurred before the organisms died would make the VEM measurements an underestimate of coseismic deformation (Figure 28). This explanation may be as unreasonable as 8 cm/yr of deformation prior to the 1992 event, however it may not be possible to differentiate between a pre-1992 interseismic signal and rapid post-1992 deformation. This illustrates how different methods of measuring coseismic deformation may capture different deformation signals.

Figure 28. Schematic depiction of how rapid postseismic deformation before intertidal organism die-off might change the VEM and uplift estimate. If any postseismic subsidence occurs before the VEM is established (organisms die or are damaged beyond recovery), then the live zone may be extended upward and the subsequent uplift measurement will be an underestimate.

Discussion of relative sea level from photo evidence

Based on a comparison of water level on rocks with a comparable time series of photos from 1992 and 2014/15 there has been no more than 20 cm of subsidence. This allows for ~10 cm of error in the actual water level due to imprecise tide calculations, swell size, wind and individual waves. It also allows ~10 cm of error in actual measurement of water level on the rocks, which has some inherent error due to photo distortion and some rock feature ambiguity.

In 1992, when these photos were taken, very low tides, as low as -1.8 ft., were occurring during the day. This was very fortunate for the 1992 VEM study, however

54 during the ~1.5 years of my study such low tides did not occur during daylight, so I was unable to recreate many of the photos at the correct tide level.

Because of the lack of 2014/15 data in this portion of the study, I cannot say that I made an accurate measurement of relative sea level change using photos, especially at

Mussel Rock and Devils Gate, where the coseismic uplift was the greatest. At these locations I expected the largest potential interseismic displacement, but the tides were not appropriate for measurement using this method. My most precise relative sea level comparisons were done south of the Mattole river mouth and at Cape Mendocino, where coseismic uplift was low and I expected little interseismic deformation.

Discussion of intertidal organism location and estimate of relative sea level

Measurement of relative sea level change since 1992 using comparison of intertidal organism locations was inconclusive due to several factors: 1) the measurements are inherently highly variable, 2) the number of comparison sites was small and 3) the current density of intertidal life in the photo locations is sparse. This study might have been more viable if the 1992 digital total station survey data were available so that population elevations could have been averaged over more data points and all compared to a fixed datum instead of being dependent on individual rocks.

I was unable to extract detailed species information from the 1992 photos and field descriptions of those species in specific photos were limited or absent because the photos had not been taken with the intent to be used in future study. Finally and most

significantly, the current intertidal communities appear to be sparse relative to those

present in 1992. This greatly hindered the study by limiting the number of species that

could be measured at each site and added uncertainty to the validity of measurements.

The current (2014/15) intertidal populations at Mussel Rock are sparse and unhealthy in

most photo locations. Mussel density is low and the patches are small. Acorn barnacles

are small and sparse compared to other reefs to the north and south. Goose barnacle

communities appear healthy at Mussel rock; they are the most populous and dense

species in the upper tidal zone. There are few large seaweed in the mid upper tidal zone,

whereas in 1992, before the uplift, the population was dense. At the time of this study in

2014 and 2015, the most populous intertidal plant, the pink coral, was starting to turn

white and were dying.

Possible reasons for the current condition of the intertidal communities may

include slow repopulation, unfavorable post-uplift conditions, currently changing relative sea level, and other uplift-independent environmental factors such as water temperature or drought-induced high sand level. Slow repopulation could be reflected in the community sparseness; if the entire population of a particular species died off in the uplift, new larvae might have to relocate from north and south, and recolonize independently. Although the populations at the reef of Mussel Rock are sparse, other reefs had large, healthy intertidal communities. The reef immediately south of Mussel

Rock (Figure 18) has thick bands of mussels and bountiful seaweeds similar to those evident in the 1992 pictures of Mussel Rock. Also to the north, at Devils gate, the seaweed is much denser than at Mussel Rock. These reefs did not experience as much

uplift as Mussel Rock, so it is possible that the entire population at Mussel rock died off in 1992, whereas locations to the north and south did not experience complete die-off and were able to recolonize more effectively.

If the entire population at a reef died off due to uplift, it could take much longer for the reef to be recolonized than if some portion of the original population survived

(Personal communication, R. Rasmussen, 2015). Larva would need to be transported to the reef from those to the north and south. Being young and weak, larva would only be able to repopulate the most favorable areas at first and could slowly expand the species distribution over many years. Another possibility would be that relative sea level is currently changing enough to have disrupted the population. However, this is not

supported by the GPS data and the mussels further west on Mussel rock reef are thick and

healthy.

It is also possible that environmental factors such as water temperature or an

invasive species could affect the communities. One such environmental factor is high

sand levels that I observed from the start of this project in late 2012 through late 2014.

Summer sand emplaced in 2012 or earlier remained on the Mussel Rock reef until late

2014, for at least a full two years there were no storms with enough wave action to

remove high summer sands. Intertidal organisms are adapted to live beneath the sand in

the summer, however, two summers and two winters of constant sand coverage may have

been too much for them to handle and caused die-off. Another indication of the influence

of beach sand is that the mussels further out on the reef are thick and healthy, only the

rocks closest to the beach are extremely sparse. Unfortunately, these areas are where the

1992 photos were taken, since the reef farther out is inaccessible.

Another possibility is the receded water has left the easternmost sites too far from

the water to support healthy mussel patches. A main assumption of this method is that if there is a nearly vertical rock face within the tide zone, organisms will live higher or lower on the rock as relative sea level is raised or lowered (Figure 29). However, if something, such as the pre-uplift wavecut platform, interferes with the wave splash when sea level is lowered (rocks are uplifted), the organisms might not re-inhabit a lower location on the same rock face, or may have a smaller patch (Figure 30).

Figure 29. Examples of how reef morphology could affect the presence of repopulation within a photo extent. The method is most effective with scenario A, but in scenarios B and C, the current living patch is completely outside the photograph area. In B and C no interseismic measurement would be possible.

Figure 30. Schematic showing how the presence of a wavecut platform between small sea stacks, irregularities or remnants of a higher platform could cause easternmost reef rocks to have smaller patches, despite being at the appropriate elevation. Although the easternmost rocks might be within the tidal zone, if the waves are very shallow when they reach the rock the water might not support a large patch.

Urchin pit discussion

Because the sea urchins are a mobile creature, they were not used in the desiccation study of 1992. I used the uplift measurement for mussel rock in the calculations, not the measurement from the 1994 photos because they are only of one location and I could not identify the exact rock from the photos. Care must be taken to measure urchin heights only in unobstructed swash channels. Urchins will live much higher on the rocks in areas were sea water is accumulated and held, such as pools on top of the terrace or in swash channels that do not fully empty. These colony elevations are not indicative of relative sea level. In the 2015 photo of figure 20, the water level is equal with the urchin colony top. This is merely a coincidence of the tide level on that day. The tide can be lower than the urchin colony top. Also coincidentally, the current algae level is fairly even with the top of the empty pits. This only shows that the algae lives within the tide that is 1.5 m above the tide level that the urchins prefer.

My measurements of relative sea level change using urchin pits were also highly

variable, particularly from one side of each swash channel to the other. However, the

measurements averaged out to within 10 cm of the coseismic uplift measurement at

Mussel Rock. This could indicate 0 – 30 cm of subsidence since 1992. There are a few

sources of uncertainty. Actual coseismic uplift at this reef may have been up to 25 cm

less than Mussel Rock to the north, although this is a liberal estimate based on the

measured die-off at Sea Lion Rock ~2 km to the south (Carver et al., 1994)(Figure 6).

The urchins based measurement is precise to within about 10 cm, due to individual

desiccation tolerances and small variations in wave splash (Personal communication, R.

Rasmussen, 2015). This uncertainty applies to both the location of current colony top

and the top of the 1992-occupied urchin pits. Despite these sources of error, my results

show that there has not been a large amount of subsidence since 1992, although it is not

precise enough to distinguish subsidence less than 30 cm.

CONCLUSIONS

In this study I measured the cumulative vertical deformation for the 23 years since

the 1992 Cape Mendocino earthquake. The maximum vertical deformation rate I have

measured is -1 mm/yr at benchmark Z1465 near Mussel Rock. The benchmarks along the

hinge between coseismic uplift and subsidence have no measureable elevation change

since 1992. Farther inland, where coseismic subsidence occurred, I have measured a

maximum of 1 mm/yr of interseismic uplift at benchmark K275. Along the coast using

photographic comparison of relative sea level and intertidal organism position, I was

unable to make precise measurements of relative sea level change. These photographic

methods indicate that there has not been more than 30 cm of subsidence at Mussel Rock,

which experienced the peak measured coseismic uplift of 1.4 m. My observations

indicate that interseismic deformation is roughly inverse of coseismic deformation,

however it is occurring at an extremely low rate. The observed deformation rate is even

lower than the interseismic rate on the rest of Cascadia, indicating that the 1992 event

was likely on a subsidiary upper plate fault and the geodetic signals of the CSZ and the

1992 earthquake fault may be interfering.

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APPENDIX

This appendix shows each intertidal organism position comparison analysis completed at Mussel Rock. The most distinct measurement points are annotated. In some photos not all comparable distribution boundaries were measured because they were not the same distance from the camera as the survey rod used for scale. If some portion of the rock was closer or farther from the camera than the scale then measurements from the photograph would not be accurate. In some locations additional measurements were made using additional photographs with the scale rod held in an appropriate location. Final measurements averaged from photo analysis are displayed in Table 3.