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Quaternary Science Reviews 113 (2015) 78e92

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Quaternary Science Reviews

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A sea-level database for the Pacific coast of central North America

Simon E. Engelhart a, *, Matteo Vacchi b, Benjamin P. Horton c, d, e, Alan R. Nelson f,Robert E. Kopp d, g

a Department of Geosciences, University of Rhode Island, Woodward Hall, Kingston, RI 02881, USAb Aix-Marseille Universit�e, CEREGE CNRS-IRD UMR 34, Europole de l'Arbois BP 80, 13545 Aix-en-Provence Cedex 4, Francec Sea Level Research, Department of Marine and Coastal Sciences, Rutgers University, New Brunswick, NJ 08901, USAd Institute of Earth, Ocean and Atmospheric Sciences, Rutgers University, New Brunswick, NJ 08901, USAe Division of Earth Sciences and Earth Observatory of Singapore, Nanyang Technological University, 639798, Singaporef Geologic Hazards Science Center, U.S. Geological Survey, Golden, CO 80401, USAg Department of Earth & Planetary Sciences and Rutgers Energy Institute, Rutgers University, Piscataway, NJ 08854, USA

a r t i c l e i n f o

Article history:Received 30 May 2014Received in revised form28 November 2014Accepted 7 December 2014Available online 15 January 2015

Keywords:Sea-level databaseCascadia subduction zonePacific North AmericaGlacial isostatic adjustmentHolocene

* Corresponding author. Tel.: þ1 401 874 2187.E-mail address: engelhart@uri.edu (S.E. Engelhart)

http://dx.doi.org/10.1016/j.quascirev.2014.12.0010277-3791/© 2014 Elsevier Ltd. All rights reserved.

a b s t r a c t

A database of published and new relative sea-level (RSL) data for the past 16 ka constrains the sea-levelhistories of the Pacific coast of central North America (southern British Columbia to central California).Our reevaluation of the stratigraphic context and radiocarbon age of sea-level indicators from geologicaland archaeological investigations yields 600 sea-level index points and 241 sea-level limiting points. Wesubdivided the database into 12 regions based on the availability of data, tectonic setting, and distancefrom the former Cordilleran ice sheet. Most index (95%) and limiting points (54%) are <7 ka; older datacome mainly from British Columbia and San Francisco Bay. The stratigraphic position of points was usedas a first-order assessment of compaction. Formerly glaciated areas show variable RSL change; wheredata are present, highstands of RSL occur immediately post-deglaciation and in the mid to late Holocene.Sites at the periphery and distant to formerly glaciated areas demonstrate a continuous rise in RSL with adecreasing rate through time due to the collapse of the peripheral forebulge and the reduction inmeltwater input during deglaciation. Late Holocene RSL change varies spatially from falling at0.7 ± 0.8 mm a�1 in southern British Columbia to rising at 1.5 ± 0.3 mm a�1 in California. The differentsea-level histories are an ongoing isostatic response to deglaciation of the Cordilleran and Laurentide IceSheets.

© 2014 Elsevier Ltd. All rights reserved.

1. Introduction

Regional databases of relative sea level (RSL)dfor example, inGreat Britain (e.g., Shennan and Horton, 2002; Bradley et al., 2011)and along the coasts of North America (e.g., Engelhart and Horton,2012; Shugar et al., 2014)dprovide a framework for developing ourunderstanding of the primary mechanisms of RSL change since theLast Glacial Maximum (~26 ka, e.g., Peltier et al., 2002). Regionaldatabases also represent a long-term baseline against which togauge changes in sea level during the 20th century (e.g., Mazzottiet al., 2008; Engelhart et al., 2009), forecasts for the 21st century(e.g., Horton et al., 2014), and the basis for identifying regional

.

variations in RSL (e.g., Shennan and Horton, 2002). Further, degla-cial RSL reconstructions are used to constrain geophysicalmodels ofglacial isostatic adjustment (GIA, e.g., Peltier et al., 2002; Mitrovica,2003; Lambeck et al., 2004; Milne and Peros, 2013). Late-Holocenedata are crucial to assess spatial variability of rates of ongoing GIA(e.g., Engelhart et al., 2009). Such information is particularlyimportant to correct sea-level measurements obtained by instru-mental methods (e.g., Church and White, 2011).

Changes in RSL are the net effect of simultaneous contributionsfrom eustatic, isostatic (glacio and hydro), tectonic and local factors,all of which have different response timescales. The relativeimportance of these factors varies in time and space along thecentral Pacific coast of North America. The greatest RSL changesince the Last Glacial Maximum was caused by the melting ofapproximately 50 million km3 of ice in land-based ice sheets,raising RSL in regions distant from themajor glaciation centers (far-

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field sites) by ~120e130 m (e.g., Lambeck, 2002; Peltier andFairbanks, 2006; Lambeck et al., 2014). This meltwater (oreustatic) contribution to far-field RSL rise during deglaciationaveraged 12 mm a�1, although peak rates potentially exceeded40 mm a�1 during “meltwater pulses” at 19 and 14.5 ka (e.g.,Deschamps et al., 2012; Lambeck et al., 2014). Empirical and glacialisostatic adjustment (GIA) modeling studies suggest a significantreduction in the meltwater contribution to RSL change at 8 ka,following which ocean volume changed by less than a few meters(Mitrovica and Milne, 2002; Bassett et al., 2005; Lambeck et al.,2014). RSL dropped by over 100 m in regions once covered by icesheets (near-field sites) as a consequence of glacial isostaticadjustment, or isostatic rebound of Earth's crust (e.g., Clague et al.,1982; Roe et al., 2013). Growth and thickening of an ice sheet re-sults in subsidence of land beneath the ice mass, which iscompensated for by an outward flow of mantle material that upliftsa peripheral bulge around the ice margin. When the ice sheet meltsand loading is diminished, land beneath themelted ice is uplifted atrates which may locally reach 50e100 mm a�1 (e.g., Shaw et al.,2002). The peripheral forebulge subsides and moves progres-sively toward the center of the diminishing load as mantle materialis redistributed (Peltier, 2004).

Along parts of the Pacific coast of central North America (Fig.1a),RSL histories are also of fundamental importance in recordingvertical tectonic land-level changes (uplift or subsidence) caused by(1) localized folding and faulting in Earth's upper crust, and (2)regional deformationdboth elastic and permanentdduring cyclesof strain accumulation and release on the megathrust fault where

Fig. 1. (a) Approximate spatial extent of the Cordilleran and Laurentide ice sheet at ~ 21.5 kacentral western North America showing active faults and plate motions (based on Wells andof modeled block movements from McCaffrey et al. (2013). Oblique subduction of the Jua35 mm a�1 causes rotation and northward movement of the Oregon Coast Range block againcompression with active faulting in the Puget Lowland of northwest Washington. The San AnSierra Nevada block. Numbered rectangles show the location of sea-level data of this papeCascadia; ORC e Oregon Coast Range; SNV e Sierra Nevada; and YAK e Yakima Fold Belt.

the Juan de Fuca plate subducts beneath North America at theCascadia subduction zone (e.g., Wells and Simpson, 2001; Wanget al., 2012; McCaffrey et al., 2013; Fig. 1b). Such histories of tec-tonic land-level change, in turn, have helped build more realisticmodels of plate-boundary and upper-plate deformation and moreaccurate assessments of the hazard posed by earthquakes on localand regional faults (e.g., Burgette et al., 2009; Wang et al., 2013;Nelson et al., 2014). As discussed for each region below, verticalrates of uplift or subsidencedintegrated over the centuries tomillennia spanned by the data points of this compilationdare onlyhigh enough to influence rates of RSL change near Holocene faultsor folds with high rates of deformation in the shallow crust (<25 kmdepth; e.g., James et al., 2009).

Local but highly spatially variable factors may influence RSLhistory. These include changes in the tidal regime (e.g., Shennanet al., 2000a; Uehara et al., 2006; Hill et al., 2011; Hall et al.,2013; Horton et al., 2013a) due to changes in local geo-morphology (e.g., Shennan et al., 2003) or to global changes in tidalamplitudes. The latter may be the result of changes in the avail-ability of sites for dissipation of tidal energy, such as Hudson Bay,that are affected by continental glaciation (e.g., Hill et al., 2011).Sediment consolidation due to compaction of pre-Holocene strata(e.g., Miller et al., 2013) through the accumulation of overlyingsediment and land drainage (e.g., Kaye and Barghoorn, 1964;T€ornqvist et al., 2008) can also produce significant errors in RSLreconstructions.

After summarizing previous Holocene sea-level studies alongthe Pacific coast of central North America, we explain how a

redrawn from Clague and James (2002) and Dyke et al. (2002). b) Simplified tectonics ofSimpson, 2001; Kelsey et al., 2012; McCaffrey et al., 2013). Gray arrows show directionn de Fuca plate beneath North America along the Cascadia subduction zone (CSZ) atst the buttress rocks of Vancouver Island (southward pointing short arrows) producingdreas Fault (SAF) accommodates most of the motion between the Pacific plate and ther grouped into regions as explained in text. B&R e Basin and Range; NCS e Northern

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consistent methodology for selecting sea-level index points im-proves on previous compilations, and outline our reasons forgrouping sea-level data into 12 regions. We reevaluated the 342sea-level data points compiled by Hutchinson (1992), added 499new sea-level data points from more recent studies, for a total of600 index points and additional 241 sea-level limiting points thatprovide maximum or minimum constraints on former sea levels.We then reconstruct and compare the RSL histories amongdifferent regions and briefly discuss some implications of the his-tory of RSL changes along the coast of central North America.

2. Previous work

The post-Last Glacial Maximum RSL history of southern BritishColumbia coasts has been investigated since the 1970s (e.g.,Mathews et al., 1970; Andrews and Retherford, 1978; Clague, 1980;Clague et al., 1982; James et al., 2000, 2005, 2009; Clague andJames, 2002; Hutchinson et al., 2004). These studies, particularlyfocused on the early to mid Holocene RSL history and on themagnitude of the sea-level lowstand, highlight the complexisostatic response of the area to downwasting and retreat of theCordilleran Ice Sheet. The timing and magnitude of the sea-levellowstand, which have important implications for understandingdifferent ice loading histories during the Last Glacial Maximum,continues to be debated (Mosher and Hewitt, 2004; Dallimore et al.,2008; James et al., 2009; Roe et al., 2013; McLaren et al., 2014).

Much sea-level data have also been collected as part of coastalinvestigations with other goals, for example, studies of the earth-quake and tsunami history of the Cascadia subduction zone, alongthe coasts of British Columbia, Washington, Oregon, and northernCalifornia (e.g., Atwater, 1992; Darienzo et al., 1994; Nelson et al.,1996; Atwater and Hemphill-Haley, 1997; Witter et el., 2003;Hawkes et al., 2011). Studies of active upper-plate faults in thePuget Lowland of northwest Washington supplied additional sea-level data (e.g., Bucknam et al., 1992; Sherrod et al., 2000; Kelseyet al., 2004, 2012). Paleo-tsunami investigations provided datapoints from Vancouver Island to Crescent City California (e.g.,Clague and Bobrowsky, 1994; Nelson et al., 2004; Williams et al.,2005; Peterson et al., 2011). Other coastal studies with sea-leveldata have focused on shoreline development and sediment trans-port (e.g., Peterson et al., 1984, 2010; Williams and Roberts, 1989;Peterson and Phipps, 1992), geo-archaeology (e.g., Friele andHutchinson, 1993; Minor and Grant, 1996; Reinhardt et al., 1996)the evolution of the Columbia River estuary (e.g., Peterson et al.,2013; Peterson, 2014), and postglacial sea-level history (e.g.,Eronen et al., 1987; Beale, 1991; Hutchinson, 1992; Long andShennan, 1998).

Post-Last Glacial Maximum sea-level databases were recentlypublished by Shugar et al. (2014) and McLaren et al. (2014) fromsouthern Oregon to Alaska, and for the central Pacific coast ofCanada, respectively. Our new database includes two additionalregions (northern and central California) and additional data fromOregon and Washington (221 sea-level index points and limitingdata). We also assess all index points more rigorously than Shugaret al. (2014) and McLaren et al. (2014), as explained below.

3. Compilation of the sea-level database

The methodology to compile a database of sea-level index andlimiting points has been defined by the International GeoscienceProgramme (IGCP) projects 61, 200, 495 and 588 (e.g., Preuss, 1979;Tooley, 1982; van de Plassche, 1982; Shennan, 1987; Gehrels andLong, 2007; Horton et al., 2009; Switzer et al., 2012; Shennanet al., 2015). A sea-level index point is a discrete reconstruction ofthe unique position of RSL in time and space (van de Plassche et al.,

2014) Where a suite of sea-level index points exists for a locality orregion, they describe changes in RSL through time and estimaterates of change. A limiting point indicates formation in a freshwateror marine environment (Engelhart et al., 2011). Therefore, recon-structed RSL must fall below freshwater limiting dates and abovemarine limiting dates (Shennan and Horton, 2002). Although thesedata cannot be used to produce valid sea-level index points, theyare extremely important in understanding and interpreting sea-level changes along the Pacific coast of central North America. Wevalidated all index points and limiting points in our database byinsuring that each sea-level indicator had (1) a location; (2) acalibrated age (in solar years); (3) an elevation at which the indi-cator was sampled, and (4) a known relationship between the in-dicator and a contemporaneous tidal level (termed the indicativemeaning; van de Plassche, 1986; Shennan et al., 2015). Sea-levelindicators are physical, biological or chemical features possessinga systematic and quantifiable relationship to elevation in the tidalframe (Shennan, 1986; van de Plassche, 1986). This methodologyimproves upon data from some previous compilations (e.g., Shugaret al., 2014) where only a generalized tidal environment (e.g.,intertidal, subtidal, supratidal) was listed for many indicators andno indicator-specific elevational range or errorswere estimated. Forexample, 342 of the data points in our database were extractedfrom the comprehensive catalogue of radiocarbon ages ofHutchinson (1992) based on data collected between Queen Char-lotte Strait and Cape Mendocino. We re-evaluated the ages used byHutchinson (1992) and added 500 additional data points frompublished and unpublished sources to our database. Many pointscompiled by Hutchinson (1992) lacked an indicative meaning,which is required for a validated index point (see section 3,Shennan, 1986; Horton et al., 2000; Shennan et al., 2015).

More than 490 of the sea-level data points are extracted fromstudies focused on dating the timing of great earthquakes andtsunamis along the Cascadia subduction zone. Almost all (>95%)such dated indicators yield maximum ages for the times of earth-quakes and tsunamis, because the ages are on either (1) detritalmaterials (wood, herb parts, seeds) from the upper few centimetersof subsided wetland soils or from the tsunami or tidal depositsimmediately above them, or plant stems or rhizomes rooted in theupper parts of the soils (e.g., Atwater and Yamaguchi, 1991; Nelson,1992; Kemp et al., 2013). If, as is commonly assumed, coseismicallysubsided wetland soils recover much of their pre-earthquakeelevation within decades of earthquakes (e.g., Long and Shennan,1998; Hughes et al., 2002; Burgette et al., 2009; Wang et al.,2013), wetland soil samples that closely predate an earthquakeprovide accurate sea-level index points. Except in a few areas, suchas near Cape Blanco (Kelsey, 1990) and Cape Mendocino (Merrittsand Bull, 1989), studies of late Pleistocene marine terraces andrates of incision in the central Coast Range suggest variable but lowrates of uplift along much of the central Pacific coast over tens ofthousands of years (Adams, 1984; Atwater, 1987; Muhs et al., 1992;Personius, 1995; Kelsey et al., 1996; Thackray, 1998; Burgette et al.,2009). Such low net rates of long-term deformation have littleaffect on the rates of postglacial sea-level rise that we report.

We excluded data points that may be significantly affected bynon-elastic tectonic uplift or subsidence, such as from sites within afew kilometers of Holocene faults and folds with high rates ofdeformation in the Puget Lowland. For example, detailed RSL dataastride the Seattle fault indicate 5e7 m of uplift at about 1.1 ka BP(Bucknam et al., 1992; Sherrod, 2001; Nelson et al., 2014). In othercases, however, sea-level data from fault studies have probably notbeen deformed at rates high enough to invalidate them for our sea-level compilation. For example, Kelsey et al. (2004) compared ratesof sea-level rise at two salt marshes on opposite sides of an oblique-slip Holocene fault on Whidbey Island. Sea-level index points

Fig. 2. (a) Plot of all validated sea-level index points. (b) Stacked histogram of validatedbasal, intercalated and isolation sea-level index points (definitions and criteria forvalidated points explained in text).

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younger than 2.5 ka on either side of the fault show a similar rate ofrise and are used in our database. Similarly, in the Humboldt Bayarea of northern California, Holocene slip rates range from~6 mm a�1 on the Little Salmon fault and 5e7 mm a�1 on the MadRiver fault (Clarke and Carver, 1992; McCrory, 2000; Hogarth et al.,2012) and large upper-plate earthquakes are documented in thearea (Valentine et al., 2012). However, the sea-level index pointsnear these faults that we used in our database (see section 5.11) didnot show anomalous elevations, as did data points near the Seattlefault in the Puget Lowland.

San Francisco Bay is the other of our regionswhere rates of upliftor subsidence on active Holocene faults may influence the elevationof sea-level data. Whereas Ryan et al. (2008) identified subsidencebetween 0.2 and 0.3 mm a�1 with maximal rates northwest of theGolden Gate area, the majority of the long-term slip on Golden Gateplatform faults caused only minor regional uplift with a maximumrate of 0.025 mm a�1. Our San Francisco Bay sea-level index pointscome from areas 70 km to the east and 60 km to the south of theGolden Gate platform faults. Slip rates on faults in these areas aremuch lower than rates on the main faults of the San Andreas sys-tem and all faults are strike-slip (Dawson and Weldon, 2013). Thus,tectonic subsidence or uplift near our index points is probably lessthan 0.02 m ka�1.

The final database includes 600 index points and 241 limitingpoints (Fig. 2a), 42 of which are unpublished. The age range of thedata spans the last 16 ka, but the majority of the index points (95%)and limiting points (54%) are younger than 7 ka (Fig. 2b). Indexpoints and limiting points older than 7 ka are generally found atsites in British Columbia, the Puget Lowland and San Francisco Bay.The complete database and the references to the original data are inAppendix 1 and Appendix 2, respectively.

3.1. Types of sea-level points

Although a variety of types of geomorphic and stratigraphicevidence can be used to define sea-level index or limiting points(e.g., van de Plassche, 1986; Horton et al., 2013b; Shennan et al.,2015), the majority of the sea-level points in our new databaseare from salt and freshwater wetlands or adjacent estuarine sedi-ment. Following Engelhart and Horton (2012), three indicativemeanings (Table 1) were utilized for sea-level indicators from salt-marsh sediment based on the zonation of vascular plant commu-nities (e.g., Nelson and Kashima, 1993) and assemblages of micro-fossils (e.g., Jennings and Nelson, 1992; Hemphill-Haley, 1995;Hawkes et al., 2010), and to a lesser extent on d13C values of bulksediment and plant macrofossils (e.g., Engelhart et al., 2013). Theindicative meaning is composed of a reference water level (e.g.,Mean High Water, MHW) that defines the relation of the indicatorto a contemporaneous tide level and is the midpoint of the indic-ative range (the elevational range occupied by a sea-level indicator(Table 1; van de Plassche, 1986; Horton et al., 2000; Shennan et al.,2015)). The indicative range for indicators from salt-marsh sedi-ment is between Highest Astronomical Tide (HAT) and Mean TideLevel (MTL). For indicators deposited in a high salt-marsh envi-ronment, the indicative range is restricted to HAT to MHW. Forthose from low salt-marsh environments, we apply an indicativerange of MHW to MTL.

If an indicator formed in a freshwater or marine environment, itis classified as a sea-level limiting point. Although limiting pointsconstrain the position of RSL less precisely than index points, theyare important in reconstructing sea-level changes because RSLmust fall below freshwater limiting points and above marinelimiting points (e.g., Shennan and Horton, 2002). Freshwaterlimiting points usually formed at an elevation above HAT, but insome cases they may have formed at elevations within the

intertidal zone due to rising groundwater tables (e.g., Jelgersma,1961; Shennan et al., 2000a; Engelhart and Horton, 2012). There-fore, a reference water level of MTL has been employed in thisanalysis. Marine limiting points formed below MTL.

In the formerly glaciated regions of British Columbia charac-terized by tens of meters of glacial isostatic uplift, isolation basinsmay provide precise sea-level index points (e.g., James et al., 2009;Roe et al., 2013). Isolation basins are natural rock depressions thatwere, at different times in their history, isolated from, or connectedto, the sea (e.g., Shennan et al., 2000b; Long et al., 2006; Hortonet al., 2013b). When RSL exceeded the elevation of a basin sill, thebasin was inundated by marine water. Conversely, when RSL waslower than the sill, the basin was isolated from the sea and fresh-water sediment was deposited (Lloyd, 2000). Isolation of a basin,caused by land uplift exceeding eustatic sea-level rise (falling RSL),is recorded in basin sediment by a change from marine to fresh-water deposits. The age of basin isolation is estimated by dating thecontact between marine and freshwater sediment. Based uponextensive work in Europe (e.g., Lloyd, 2000; Shennan et al., 2000b;Long et al., 2006), we conservatively estimate an indicative rangefor the transition frommarine to freshwater sediment to be HAT toMTL. The height of RSL at the time of isolation is precisely con-strained by the elevation (with respect to modern sea level) of thebasin sill and not by the elevation of the contact between marineand freshwater sediment (Horton et al., 2013b).

Table 1Summary of the indicative meanings used to estimate the relative elevation of the sea-level index points and limiting points for the database. HAT e Highest AstronomicalTide; MHHW e Mean Higher High Water; MHW e Mean High Water; and MTL e Mean Tide Level.

Sample type Evidence Reference water level Indicative range

Index PointsHigh marsh environment High marsh plant macrofossils (e.g. Eicher, 1987; Nelson, 1992; Nelson and Kashima,

1993). Foraminiferal and diatom assemblages dominated by high marsh taxa (e.g.,Hemphill-Haley, 1995; Nelson et al., 2008; Hawkes et al., 2010)

(HATeMHW)/2 HATeMHW

Low marsh environment Low marsh plant macrofossils low marsh plant macrofossils. Foraminiferal and diatomsassemblages dominated by low marsh taxa (e.g., Scott et al., 1996; Hawkes et al., 2010).

(MHWeMTL)/2 MHWeMTL

Undifferentiated saltmarsh environment

Author listing of unnamed salt marsh plant macrofossils or identification only to genuslevel (e.g., Jennings and Nelson, 1992; Nelson, 1992; Nelson and Kashima, 1993)Foraminiferal and diatom assemblages dominated by high and low marsh taxa (e.g.,Shennan et al., 1996, 1998; Jennings and Nelson, 1992).

(HATeMTL)/2 HATeMTL

Isolation Basin Sediments recording the switch between freshwater and marine sediments that aresupported by changing diatoms assemblages (e.g., James et al.., 2009; Roe et al., 2013).

(HATeMTL)/2 HATeMTL

Limiting PointsMarine limiting Identifiable in-situ marine shells (e.g., Atwater et al., 1977; James et al., 2005) Calcareous

foraminiferal assemblages (e.g., Jennings and Nelson, 1992; Scott et al., 1996) in clasticsediment may be supported by diatom assemblages dominated by marine taxa(Shennan et al., 1998). Isolation basin sediments with marine diatoms (e.g. Hutchinsonet al., 2004) and marine shell assemblages (e.g. Hutchinson et al., 2004; Mosher andHewitt, 2004).

MTL Below MTL

Freshwater limiting In-situ tree stumps (e.g., Atwater and Hemphill-Haley, 1997) Peat that does not meetthe above requirements to be classified as an index point (e.g., Shennan and Horton,2002). Peat with freshwater diatoms. Freshwater isolation basin sediments (lacustrinegyttja) dominated by freshwater diatoms and freshwater shells assemblages (e.g.Hutchinson et al., 2004).

MTL Above MTL

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We also evaluated index and limiting points based on stratig-raphy described in archaeological investigations in BritishColumbia, Washington, Oregon, and California. Evidence of coastalhabitation (e.g., midden deposits) from Washington, Oregon, andnorthern California indicate occupation of most sites within the last2 ka (Woodward et al., 1990; Minor and Grant, 1996). The nativepeoples of southern British Columbia shared an economy heavilybased on maritime, intertidal and riverine resources (Hutchinsonand McMillan, 1997) and archaeological investigations in this re-gion span the last 6000 years (Grier et al., 2009). Sites of formervillages are marked by extensive midden deposits, composedlargely of discarded shellfish remains and other domestic refuse.Villages consisted of rows of plank-clad houses near the water,usually only a few meters above sea level. Due to the many un-certain factors in relating midden deposits to a contemporary tidallevel, archaeological data are interpreted as freshwater limitingpoints with a reference water level of MTL.

3.2. Elevation of former sea levels

RSL for each dated indicator is estimated using the followingequation (Shennan and Horton, 2002):

RSLi ¼ Ei � RWLi (1)

where Ei and RWLi are, respectively, the elevation and the referencewater level of indicator i, both expressed relative to the same da-tum, MTL in our analysis.

The associated total vertical uncertainty is expressed by thefollowing equation (Shennan and Horton, 2002):

E ¼�e21 þ e22 þ e23 þ e2n…

�1=2(2)

where e1, en represent the sources of error for each index pointincluding the indicative range. The additional errors are thoseassociated with calculating the indicator elevation. Such errors canbe as small as ±0.05 m for the errors in high precision surveying(e.g., Shennan, 1986), or greater than ±0.5 m when the elevation is

estimated from the tidal environment fromwhich the indicator wascollected (e.g., salt marsh; Engelhart and Horton, 2012). Sea-levelindicators without elevational data were rejected. We included anerror of ±0.1 m to account for the stability of benchmarks (NationalGeodetic Survey classification, Horton et al., 2009). Indicatorthickness is also incorporated into the error term. For some olderbulk peat samples, this may be as large as ±0.25 m (e.g., Vick, 1988;Valentine, 1992). We also calculated the error in elevation due tothe angle of boreholes as a function of indicator depth, taken in thisstudy as ±1% (T€ornqvist et al., 2008).

We do not include an error term for potential paleotidal rangechange (e.g., Uehara et al., 2006; Hill et al., 2011; Horton et al.,2013a) or sediment compaction (e.g., Edwards, 2006; Horton andShennan, 2009). Sediment compaction results in a lowering ofthe index point relative to the initial depositional elevation and willresult in an overestimation of RSL rise. Such processes have beennoted, for example, for indicators from buried former marshes atthe Salmon River estuary (Nelson et al., 2004) in comparison tothose from buried former marshes at other Oregon estuaries (Longand Shennan, 1998). We make a first order assessment of sedimentcompaction by subdividing index points into basal and intercalatedcategories (Shennan and Horton, 2002; Horton and Shennan,2009). Basal index points are from sediment that directly overliesincompressible substrate (Jelgersma, 1961). Intercalated indexpoints are derived from beds of low-density, organic-rich sedimentwithin a sequence of higher density, clastic units (Shennan andHorton, 2002) and are, therefore, particularly susceptible tocompaction (Jelgersma, 1961). Where stratigraphic informationwas unavailable for an index point we conservatively interpreted itas intercalated. Index points from isolation basins are compactionfree as the height of the bedrock sill, and not the height of thesampled sediment, controls the position of former RSL.

3.3. Age of sea-level indicators

All ages in our database are radiocarbon (14C) ages on organicmaterial from salt and freshwater marshes or marine shells. In-dicators were dated using different methods of radiocarbon

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analysis, including Gas Scintillation Counting (GSC), Liquid Scintil-lation Counting (LSC), and Accelerator Mass Spectrometry (AMS).All indicators were calibrated using CALIB 7.0 (Stuiver et al., 2013)with a 2s error. We employed the IntCal13 and Marine13 (Reimeret al., 2013) datasets for terrestrial and marine indicators, respec-tively. For shells we used carbon reservoir corrections from theMarine Reservoir Database (Reimer and Reimer, 2001) or fromother published values for the region (Appendix 1). All index pointsare presented as calibrated years before present (ka), where year0 is 1950 CE. Basal gyttja ages can be affected by the introduction ofold carbon from underlying marine sediments (Hutchinson et al.,2004; James et al., 2009). A correction factor of �625 ± 60 yearswas applied to these ages (e.g., Hutchinson et al., 2004; James et al.,2009; Roe et al., 2013). Twentieth-century ages reported by theGeological Survey of Canada (GSC) radiocarbon lab are corrected toconventional ages following Stuiver and Polach (1977). A concernwith radiocarbon ages analyzed prior to the mid-1970s is thecorrection used for isotopic fractionation (e.g., Hijma et al., 2015;T€ornqvist et al., 2015). In the database, the majority of ages wereanalyzed after 1980 and, therefore, most are not subject to thispotential error. For the 37 ages that are affected, we followed theprocedure of Hijma et al. (2015) to correct for isotopic fractionation.However, the general absence of C4 plant material within saltmarshes on the Pacific coast of North America (e.g., Engelhart et al.,2013) minimizes this potential error in ages, as most datedmaterialprobably had values close to �25‰ relative to the VPDB standard.

3.4. Statistical framework to assess spatial variability of lateHolocene relative sea-level change

We assessed the spatial variability of GIA and tectonicallyinduced land-level changes using an empirical-Bayesian spatio-temporal statistical model for late Holocene (<4 ka) sea-levelchange in which RSL at each site is assumed to change linearlyover time at a ratem(x). The fieldm(x) has a Gaussian process prior(Rasmussen and Williams, 2006) with an exponential covariancefunction:

m xð Þ � GP 0; s2s exp�r x; x0ð Þ

l

� �� �(3)

where s2s is the prior variance, rðx;x0 Þ the angular distance between

x and x′, and l is a length scale. The index point observations yi aremodeled (limiting data are not modeled):

yi ¼ m xið Þ bti � t0 þ εti

� �þ ε

yi þwi þ yigi þ y0 xið Þ: (4)

The reference time t0 is set to 1970 CE. The term bti is the meanestimated age of the observation, which has normally distributederror εti with variance specified for each point in the database. Thenoise term ε

yi is normally distributed and uncorrelated in space and

time, with variance specified for each data point in the database.The additional noise term wi is also normally distributed and un-correlated in space and time, with prior variance s2w; this term al-lows the model to increase its noise estimate to improve the fit tothe data. We allow for compaction in this model through the thirderror term, gi which is normally distributed with prior variance s2c .This term is multiplied by yi, so deeper samples, with more op-portunity for compaction, have a larger error. The term y0(xi),which has a prior variance of s20 and is spatially uncorrelated, al-lows for offsets in the datums at different sites.

We employ simulated annealing followed by local optimizationto find the hyperparameters that maximize the likelihood of themodel, conditional upon all the index points in the data set:ss ¼ 1.3 mm/y, l ¼ 29.3�, sw ¼ 0.4 mm, sc ¼ 0.0 and s0 ¼ 380 mm

(Note that, with the exception of the datum offset, the additionalnoise terms are essentially unneeded by the maximum-likelihoodmodel, implying the vertical error specified in the database isadequate.) The rate estimates provided in this analysis are based onm(x) as estimated with the maximum likelihood hyperparameters.

4. Sub-regions of sea-level data

4.1. Database subdivision

We compiled sea-level data from selected parts of the coasts ofnorthern and central California, Oregon, Washington, and southernBritish Columbia (Fig. 1b). Hutchinson (1992) subdivided his sea-level database into sub-regions according to latitude and longi-tude. Instead, we divided our database into sub-regions based on(1) regional tectonic setting, (2) distance from the former Cordil-leran ice sheet, (3) proximity to other sea-level data points(commonly a function of local geomorphology, such as estuarinelowlands between coastal headlands), and (4) location relative toHolocene faults and folds.

In simplified form, the regional-scale tectonics of westernmostcentral North America consists of shear interaction between thePacific and North American plates south of the Mendocino TripleJunction, and oblique convergence between North America and theJuan de Fuca plate north of the Mendocino Triple Junction (Fig. 1b).South of the triple junction North America deforms through strike-slip faulting, largely along the San Andreas Fault system, whereas inthe Cascadia forearc, north of the triple junction, plate subductionand the rotation of crustal blocks appear to accommodate therelative motions of tectonic plates (Wells and Simpson, 2001;McCaffrey et al., 2013). As a consequence of oblique subduction,Cenozoic geology maps show the northward migration and clock-wise rotation of the Cascadia forearc relative to North America(Wells and McCaffrey, 2013) and the most recent GPS data confirmthat the rotation continues east of the forearc into the Basin andRange (McCaffrey et al., 2013). In northern Washington, motions ofthe Oregon coast block compress the Puget Lowland against thebuttress of southwestern British Columbia. The current crustalshortening in the lowland of 4e5 mm a�1 is mostly accommodatedby slip on eastewest, Holocene thrust faults a few tens of kilome-ters long that bound basins and uplifts (Fig. 1b, Blakely et al., 2002;Mazzotti et al., 2003; McCaffrey et al., 2013). In the San FranciscoBay region of central California, about 70e80% of the 45 mm a�1 oflateral slip between the Pacific and North America plates occursalong the complex faults of the San Andreas system (Grove et al.,2010; Tong et al., 2013).

Vertical land-level changes (uplift and subsidence) along coastsof Cascadia and San Francisco Bay are largely the result of inter-seismic strain accumulation and release on faults of these regionsduring many earthquake cycles (Wang, 2007; Wang et al., 2012).The classic models of the earthquake deformation cycle invokeelastic deformation, all of which is assumed to be recoverable(Yeats et al., 1997; Scholtz, 2002). But at a particular site on acoastline, other factors such as RSL rise or fall, the amount of slip(proportional to earthquakemagnitude) and how it was distributedon the fault plane (location of asperities, updip, downdip), theposition of the site relative to the slip distribution, and viscoelasticeffects over years to many decades following earthquakes, allmodify these simple models of deformation (Wang, 2007; Wanget al., 2012; Nelson, 2013). At Cascadia, the megathrust iscurrently fully or close to fully locked, leading to margin-normalshortening and uplift of the forearc (Miller et al., 2001; Burgetteet al., 2009). This leads to significant variability in ongoing crustalstrain rates, especially in the northern portion of the subductionzone (McCaffrey et al., 2007). Vertical motions recorded by

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instrumental records (e.g., Mazzotti et al., 2008; Komar et al., 2011)represent ongoing interseismic deformation.

The complex geometry and interactions among faults beneaththe wetlands of the San Francisco Bay region makes their post-glacial vertical slip rates difficult to discern. In most cases, verticalrates are probably at least an order of magnitude lower than lateralrates, that is, <0.5 mm a�1 (e.g., Grove et al., 2010). Late Pleistoceneuplift rates measured from an 80-ka terrace along a fault northwestof San Francisco are 0.2e1.0 mm a�1 (Grove et al., 2010), whereasRyan et al. (2008) measured subsidence rate of 0.2e0.3 mm a�1 ona block bounding the San Andreas Fault system west of San Fran-cisco Bay. Marsh elevations in south San Francisco Bay have keptpace with tens of centimeters of sediment compaction and subsi-dence due to groundwater withdrawal, and to a much lesser extentto sea-level rise and tectonic uplift or subsidence, over at least thepast half century at rates as high as 40 mm a�1 (Patrick andDeLaune, 1990).

The Cordilleran ice sheet formed over British Columbia, south-ern Yukon Territory and southern Alaska (Fig. 1a; Clague and James,2002). Its glacial history, including the timing of the Last GlacialMaximum, differs from that of the Laurentide ice-sheet (Gregoireet al., 2012). At 21.4 ka, ice cover in North America was at or nearits maximum everywhere except along the southern Cordilleran icemargin (Dyke, 2004). In particular, the southwestern portion of theCordilleran ice-sheet continued to build to its maximum extentuntil ~17.2 ka (Clague and James, 2002) with expansion towards thecoast along the Strait of Georgia, as expressed by the Puget and Juande Fuca lobes (Clague,1983; James et al., 2009; Fig.1a). Deglaciationproceeded very rapidly along the southwestern Cordilleran icemargin; Juan de Fuca Strait and the Strait of Georgia were free of iceby ~16.7 ka and ~13.8 ka, respectively (Barrie and Conway, 2002;Mosher and Hewitt, 2004).

We subdivided the formerly glaciated areas into seven regions(#1 to #7, Fig. 1b). The Queen Charlotte Strait (#1) is the north-ernmost region of the database. It is located between two platetriple junctions that mark the limits of the Explorer microplate(Audet et al., 2008; Fig. 1a). GPS-derived velocity fields indicate asignificant deviation from the typical pattern of elastic strain ratesresulting from locking on the subduction zone (McCaffrey et al.,2007). South of 50�N, GPS vertical velocities show an eastewestgradient with sites on the west coast indicating uplift at1e4 mm a�1 and sites more inland indicating near-zero verticalmotion or slight subsidence (Mazzotti et al., 2008). Therefore, wesubdivided the database on the basis of these modern observations.On western Vancouver Island (#2), GPS vertical velocities showuplift rates ranging between ~2.6 and ~3.9 mm a�1. They decreasedown to ~2 mm a�1 on eastern Vancouver Island (#3), to~1.4 mm a�1 in southern Vancouver Island (#5) and to ~0.7 mm a�1

in southeastern Georgia Strait (#4). The northwestern Washingtoncoast (#6) and Puget Sound (#7) represent the southernmost limitof the former ice sheet (Clague, 1983; Clague and James, 2002;James et al., 2009). In this area, a permanent northesouth short-ening is superimposed on the transient subduction signal(Hyndman et al., 2003; Mazzotti et al., 2003) and GPS vertical ve-locities showhigh variability. They range from ~3.9mm a�1 of uplifton the northwestern Washington coast to zero vertical motion orslight subsidence in the Puget Lowland area (Mazzotti et al., 2008).

The remaining regions (#8 to #12, Fig 1b) were not covered bythe Cordilleran ice sheet during theWisconsinan glaciation (Clague,1983; Hutchinson et al., 2004). The southern Washington region(#8) marks the transition between the region of northesouthshortening and the region of northwards motion of the Oregonblock. Here, GPS vertical velocities show the highest uplift withrates >2.7 mm a�1 (Mazzotti et al., 2008). The pattern, due to thelocking of the subduction plate boundary along the southern

portion of the Cascadia margin, results in variability in uplift ratesalong the Oregon and northern California coast (Burgette et al.,2009; Schmalzle et al., 2014). Therefore, we subdivide the Oregoncoast into two regions: the northern Oregon coast (#9), where GPSvelocities indicate zero motion to slight subsidence, and the centralOregon coast (#10) where uplift rates of ~1.1mm a�1 were recorded(Mazzotti et al., 2008). The northern California region (#11) islocated roughly 100 km north of the Cape Mendocino triple junc-tion, which marks the southern limit of the Cascadia subductionzone. Deformation in this region is dominated by Gor-daePacificeNorth America oblique plate convergence, with inter-nal deformation of the Gorda plate, and lateral deformation alongthe Mendocino transform fault (Williams et al., 2006). CentralCalifornia (#12) is the southernmost region of the database and isinfluenced primarily by strike-slip deformation along the SanAndreas Fault system and related faulted systems to the south andeast.

5. Sub-region sea-level histories

5.1. Queen Charlotte Strait (#1)

The RSL history is composed of 4 index points and 13 limitingpoints (Fig. 3, #1) from the isolation basins of Seymour and BelizeInlets and offshore cores from Cook Bank (place names in Appendix1). At ~13.0 ka a single index point places RSL at 1.3± 0.9m. At ~12 kaRSL dropped dramatically to ~�95 m. There is an absence of indexpoints during the early Holocene, although freshwater limitingpoints suggest RSLwas below3.6m from~12.0 to ~8.3 ka. At ~8.0 ka,one index point places RSL above present at ~1.3 m. The late Holo-cene record suggests RSL was at ~1.3 m from ~4.1 to ~1.5 ka.

5.2. Western Vancouver Island (#2)

Western Vancouver Island data consist of 39 index points and 30limiting points (Fig. 3, #2). Salt-marsh sediment provided indexpoints from Tofino, Hesquiat Peninsula, Barkley Sound, NootkaSound, Esperanza Inlet and Alberni Inlet, and isolation basinsyielded index points and limiting dates from Effingham Inlet andMeares Island (place names in Appendix 1). Data from Lateglacialisolation basins record a very rapid RSL fall. The oldest marinelimiting point indicates RSL was above 25 m at ~15.5 ka. Twolimiting points constrain RSL to ~4 m MTL at ~14.0 ka, whereasindex points indicate a sea level of ~�46.9 m between ~13.0 and~12.3 ka. Freshwater limiting points constrain RSL to below�9m at~9.0 ka. Younger intercalated index points and freshwater limitingpoints indicate RSL was 2.5 ± 1.1 m MTL at ~6.8 ka, suggesting arapid rise to a highstand during the early Holocene. Further inter-calated index points suggest sea level was at 4.4 ± 0.7 m MTL at~5.6 ka, and then fell to 3.8 ± 0.7 m MTL at ~4.5 ka. At ~2.5 ka, oneintercalated index point and two limiting points constrain RSL to2 ± 1.1 m MTL. Younger index points indicate RSL remained abovepresent MTL until ~0.7 ka.

5.3. Northwestern Georgia Strait (#3)

The RSL history of eastern Vancouver Island is constrained by 12index points and 51 limiting points (Fig. 3, #3). Index points werefrom isolation basins on Lasqueti, Quadra and Cortes Islands andsalt marshes near Qualicum and Courtenay (eastern VancouverIsland; place names in Appendix 1). Limiting samples were fromarchaeological excavations near Comox and Nanaimo, from isola-tion basins on Lasqueti, Quadra and Cortes Islands, and from thecoast of eastern Vancouver Island. Marine and freshwater limitingpoints constrain RSL to ~150.0 m at ~13.9 ka. Multiple terrestrial

Fig. 3. Sea-level index points for regions (brown rectangles) #1 to #7 (Fig. 1b) along the Pacific coast of North America. Black dots on map show sampling sites. Index points (thosefrom isolation basins are black, basal points are blue, and intercalated points are red) are plotted as calibrated age BP against elevation relative to present sea level (m). Limitingpoints are plotted as freshwater (maximum; green) or marine (minimum; brown) horizontal lines. Dimensions of boxes and lines for each point based on 2s elevation and ageerrors. BC e British Columbia, Canada; WA eWashington, USA; CSZ e Cascadia Subduction Zone; Qcs e Queen Charlotte Strait; Gs e Georgia Strait; Jfs e Juan de Fuca Strait; Ps ePuget Sound.

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and marine limiting points then show a rapid sea-level fall to anindex point with a RSL of �0.9 ± 1.1 m at ~12.0 ka. Marine limitingpoints constrain RSL to above ~11.0 m at ~11.0 ka and above �25 at~11.0 ka. From ~9.8 ka to ~9.0 ka, limiting points constrain RSL tobetween �20 and 0 m. At ~8.1 ka, one marine limiting point sug-gests RSL was above present. Intercalated index points indicate RSLwas at 1.0 m ± 1.1 at ~6.5 ka and 2.5 ± 1.1 m at ~3.2 ka. Youngerindex points suggest RSL rose to near present at ~1.6 ka.

5.4. Southeastern Georgia Strait (#4)

The RSL database for southeastern Georgia Strait consists of 23index points and 42 limiting points (Fig. 3, #4). Archaeological ex-cavations and salt and freshwater marshes from the Fraser Riverdelta, Howe Sound, Padilla Bay and Bellingham Bay provided thedata (place names in Appendix 1). Additional limiting points werefrom the glaciomarine deposits in the Fraser River Valley and from

freshwater marshes and archaeological excavations in the GulfIslands and Skagit Bay. Marine and terrestrial limiting points indi-cate a rapid fall in RSL at the transition from the Lateglacial period tothe early Holocene. A single marine limiting point suggests that RSLwas above 154 m at ~14.6 ka. At ~13.3 ka, RSL was between 66 and58m, and between ~12.0 and ~10.0 ka limiting points constrain RSLto between 12mand�33m. At ~9.2 ka, RSLwas below�11m. Then,RSL rapidly rose during the mid-Holocene from �10.8 ± 1.0 m at~8.1 ka to �3.9 ± 1.0 m at ~5.9 ka. RSL rise apparently slowed in thelate Holocene. Intercalated index points of late Holocene age showscatter, likely illustrating the influence of compaction. The youngestpoints indicate that RSL rose to near present at ~0.8 ka.

5.5. Southern Vancouver Island (#5)

RSL history for southern Vancouver Island is derived from 10index points and 50 limiting points (Fig. 3, #5). Isolation basins

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(near Victoria and along eastern Juan de Fuca Strait), salt andfreshwater marshes and archaeological excavations (MontagueHarbor, Sooke Basin and Saanitch Inlet; place names in Appendix 1)provided the data. An index point shows RSL at 75 ± 2.1 m at~14.5 ka before rapidly falling to�2.1± 1.3mMTL at ~13.6 ka. In theearly to mid Holocene (~12.0 to ~6.5 ka), RSL was between �42.5and �1.4 m, and during the mid to late Holocene (~6.0 to ~3.3 ka),RSL was between�3.2 and 0m. Two intercalated index points placeRSL at �1.9 ± 0.9 m MTL at ~3.6 ka and at �0.9 ± 0.9 at ~1.9 ka.Younger limiting points constrain RSL between 0 and �5 m MTL.

5.6. Northern Washington coast (#6)

The RSL history for northern Washington coast consists of 14index points and 3 limiting points (Fig. 3, #6) and is restricted to thelast ~2.5 ka. Data came from the salt and freshwater marshes of theWaatch River, Crescent Bay and Neah Bay (place names in Appendix1). The oldest index point at ~2.4 ka shows that RSL wasat �0.3 ± 0.9 m MTL. In the last ~1.0 ka, index points show scatter.Basal index points indicate RSL was above present (0.5 ± 0.9 mMTL) at ~0.6 ka.

5.7. Puget Sound (#7)

RSL history for the Puget Lowland region is restricted to theLateglacial period and to the mid- and late Holocene (Fig. 3, #7).The database includes 35 index points and 9 limiting points fromthe marshes of Whidbey Island, Hood Canal, Quilcene Bay, Dis-covery Bay, the Nisqually delta, Cultus Bay, Bainbridge Island, andarchaeological excavations east of Seattle (place names in Appendix1). A marine limiting point indicates RSL was above 25 m at~13.6 ka. The oldest index point constrains RSL to �5.4 ± 1.2 m at~5.8 ka. Two basal index points place RSL at �4.5 ± 0.8 m MTL at~4.6 ka. RSL rose to �2± 0.8 m MTL at ~2.5 ka. From 2.5 ka topresent RSL rose more slowly: intercalated index points indicatethat RSL rose to within 1 m of modern MSL from 0.8 ka to present.

5.8. Southern Washington coast (#8)

Our RSL history for the southern Washington coast spans theearly Holocene to the present. The data consist of 128 index pointsand 4 limiting points from salt marshes and estuarine sediments ofWillapa Bay, Grays Harbor and the Columbia River (Fig. 4, #8). Theoldest index point places RSL at �53 ± 1.3 m at ~10.0 ka. Then, RSLrapidly rose to�21 ± 1.2 m at ~8.3 ka and to �13 ± 1.2 m at ~6.5 ka.There is considerable scatter in the mid to late Holocene interca-lated index points, likely reflecting the influence of compaction. Inthe mid Holocene, RSL rose from �4.9 ± 0.4 at ~5.1 kato �4.5 ± 1.0 m at ~4.6 ka. RSL in the late Holocene rosefrom �2.3 ± 0.4 m at ~3.6 ka to �1.2 ± 0.5 m at ~1.9 ka.

5.9. Northern Oregon coast (#9)

The Holocene RSL history of the northern Oregon coast (Fig. 4,#9) includes 110 index points and 13 limiting points from salt andfreshwater marshes, estuarine sediments, and archaeological in-vestigations of Tillamook Bay, Clatsop Plain, Salmon River estuary,Netarts Bay and Nehalem River estuary (Fig. 4, #9; place names inAppendix 1). In the early to mid-Holocene marine limiting dataindicate RSL was above �28.7 m MTL at ~9.3 ka and above �19.3 mMTL at ~8.8 ka. Two intercalated index points suggest RSL wasat �11.60 ± 1.0 mMTL at ~6.5 ka and at �10.6 ± 1 mMTL at ~6.0 ka.One basal peat point places RSL at �7.3 ± 1.0 m MTL at ~4.7 ka.There is a considerable scatter in the late Holocene intercalated

index points. However, one basal index point suggests RSL wasat �2± 1 m MTL at ~2.0 ka.

5.10. Central Oregon coast (#10)

The Holocene RSL history of the central Oregon coast is con-strained by 110 index points and 20 limiting points from the saltand freshwatermarshes, and estuarine sediments of the Alsea Riverestuary, Yaquina Bay, Coos Bay, Coquille River, Siuslaw River, andarchaeological investigations at Tahkenitch Lake (Fig. 4, #10; placenames in Appendix 1). The oldest freshwater limiting point placesRSL below �54.5 mMTL at ~11.5 ka. Marine limiting points suggestRSL was above �30.0 m MTL at ~9.5 ka, �18.9 m MTL at ~8.3 ka,and �12.6 mMTL at ~7.4 ka. Freshwater and marine limiting pointsconstrain RSL to between �7 m and �10 m MTL at ~6.5 ka. Onebasal index point suggests that RSLwas�7.9 ± 0.5mMTL at ~6.6 ka,and another that it rose to �5.3 m ± 1.1 m MTL at ~4.6 ka. A thirdshows that RSL was at �3.9 ± 0.5 m MTL at ~4.0 ka.

5.11. Northern California coast (#11)

RSL history in northern California coast is restricted to the lateHolocene. The database is composed of 45 index points and 6limiting point from the salt and freshwater marshes of HumboldtBay and near Crescent City (Fig. 4, #11). The oldest basal index pointat ~4.0 ka places RSL at �6.3 ± 1.0 m MTL, and anther show that itrose to �6.1 ± 0.1 m MTL at ~3.5 ka. There is a scatter in the indexpoints between ~4.0 and ~2.0 ka. Most of the scatter is likely relatedto compaction rather than to uplift or subsidence adjacent to theLittle Salmon and Mad River fault zones. Scatter in the index pointsdecreases in the last ~2.0 ka. RSL was at�2.0 ± 0.1 mMTL at ~1.5 kaand then rose to �1.5 ± 0.1 m MTL at ~1.1 ka.

5.12. Central California coast (# 12)

The Holocene RSL database for the central California coast iscomposed of 70 index points from the salt marshes of the Sacra-mento River Delta and southern San Francisco Bay (Fig. 4, #12). RSLwas �40.1 ± 1.7 m at ~10.4 ka and rose rapidly to �19.3 ± 1.1 m at~9.3 ka. The rise slowed in the mid Holocene when RSL wasat�9.8 ± 1.1 mMTL at ~7.8 ka and at�8.2 m ± 0.6 m at ~6.5 ka. LateHolocene index points show a continuous rise during the last~3.5 ka.

6. Discussion

6.1. Deglacial relative sea-level history of the Pacific coast of centralNorth America

Analysis of the RSL histories from 16 ka to present highlightssignificant differences between regions formerly covered by theCordilleran ice-sheet and those located south of the ice margin. Forthe regions that were formerly covered by the Cordilleran ice sheet(Fig. 3, regions #1 to #7), RSL histories are strongly controlled bythe ice unloading history, resulting in variable RSL histories due tothe complex interplay of eustatic sea-level change and isostaticadjustment of the crust. This is typically expressed as a RSL fallfollowing deglaciation due to isostatic uplift, followed by slow ratesof RSL rise during the mid to late Holocene. Regions with recordsextending beyond the Holocene (Fig. 3, #1, #2, #3, #4; #5, #7) showthe effects of GIA due to Cordilleran ice sheet retreat (e.g., Clague,1980; Dethier et al., 1995; Clague and James, 2002) with RSL asmuch as 154m above present at 14 ka. Significant differences in RSL(>45 m) exist at ~12 ka between Queen Charlotte Strait (#1) andwestern Vancouver Island (#2), which likely reflect the effects of

Fig. 4. Sea-level index points for regions (brown rectangles) #8 to #12 (Fig. 1b) along the Pacific coast of the United States. Black dots on map show sampling sites. Index points(those from isolation basins are black, basal points are blue, and intercalated points are red) are plotted as calibrated age BP against elevation relative to present sea level (m).Limiting points are plotted as freshwater (maximum; green) or marine (minimum; brown) horizontal lines. Dimensions of boxes and lines for each point based on 2s elevation andage errors. BC e British Columbia, Canada; WA eWashington; OR e Oregon; CA e California; CSZ e Cascadia Subduction Zone; SAF e San Andreas Fault; Gh e Grays Harbor, Nb e

Netarts Bay; Ab e Alsea Bay; Cb e Coos Bay; Cc e Crescent City; Hb e Humboldt Bay; Sfb e San Francisco Bay.

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differing ice load history. Clague and James (2002) show thatwestern Vancouver Island (#2) sites were deglaciated between ~18and ~14 ka, earlier than sites in Queen Charlotte Strait (#1) thatwere still ice covered at ~14 ka. Both in northwestern Georgia Strait(#3) and southern Vancouver Island (#5) RSL falls below presentbetween ~14 and ~12 ka before rising until 7 ka due to decreasingisostatic uplift and continuing high rates of eustatic sea-level rise(Lambeck et al., 2014).

Our database provides insight into the variability in the re-ported depth of the sea-level lowstand across the region followingdeglaciation in British Columbia (James et al., 2009). In southernVancouver Island (#5), James et al. (2009) placed the lowstandat �30 ± 5 m at ~11 ka. This depth is ~20 m shallower than thelowstand depth previously proposed (Linden and Schurer, 1988;Mosher and Hewitt, 2004). The variability in lowstand elevationis due to different interpretations of marine terraces identified byseismic data that suggest differences in ice loading, and hence GIAhistory. Along the western coast of Vancouver Island (#2),

isolation basin-derived index points from Effingham Inlet suggesta lowstand of ~�47 m at ~13 ka. This is a deeper lowstand thanfound in eastern Vancouver Island (#3) and southeastern GeorgiaStrait (#4), where marine limiting points constrain the lowstand toabove ~�11 and ~�33 m from ~12 to ~10 ka. James et al. (2009)demonstrated that these differences are expected due to differ-ential isostatic effects. Less definite conclusions can be drawnabout the characteristics of the lowstand along southern Van-couver Island (#5). Marine limiting points indicate RSL wasabove �42 m at ~14.5 ka and above �32 m at ~9 ka, but these datado not preclude a higher or lower RSL between these ages. Thelack of early Holocene data precludes identifying the elevations oflowstands along the northern Washington and Puget Lowlandcoasts (#6, #7).

Four of the formerly glaciated regions had mid-Holocene high-stands, which varied in elevation and temporally. In Queen Char-lotte Strait (#1), RSL rose to ~1.3 m at ~8.0 ka. At western VancouverIsland (#2), RSL rose above present between ~7.2 and ~6.4 ka before

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reaching a maximum of ~4.8 m at ~5.6 ka, as indicated by anintercalated index point on Vargas Island. At northwestern GeorgiaStrait (#3), marine and terrestrial limiting data suggest RSL wasabove present between ~9 and ~8 ka, and an intercalated indexpoint in Rosewall Creek shows it reached a maximum of ~2.5 m at~3.3 ka. These highstands are the result of the complex interplay ofglobal meltwater input and the locally variable GIA. During theearly Holocene, meltwater input outpaced the exponentiallydeclining isostatic uplift rate (e.g., T€ornqvist and Hijma, 2012).However, with the decrease in meltwater input starting at ~8 ka(Lambeck and Chappell, 2001; Lambeck et al., 2014) due to the finalmelting of the Laurentide Ice Sheet (e.g., Carlson et al., 2008),isostasy became the dominant control on RSL history. This resultedfirst in a stable highstand above present sea level before RSL fellwith a further decrease in meltwater input after ~4 ka (e.g., Peltierand Tushingham, 1991; Fleming et al., 1998; Milne et al., 2005;Engelhart et al., 2009). Similar RSL histories have been describedfrom regions towards the periphery of ice sheet margins innortheastern Scotland (e.g., Shennan et al., 2002), but historiesdiffer at sites in similar ice margin proximal settings in Maine (e.g.,Engelhart and Horton, 2012) where no Holocene RSL highstand isrecorded. The absence of a highstand with RSL rising continuouslysince ~6 ka at sites along southeastern Georgia Strait (#4) mostlikely represents persistent GIA-related subsidence during the midand late Holocene. A lack of index points in southern VancouverIsland (#5) prevents a more detailed reconstruction of mid-Holocene RSL history.

Sites from south of the Cordilleran ice margin (Fig. 4, #8 to #12)demonstrate a continuous rise of RSL throughout the Holocenedue to the collapse of a proglacial forebulge (e.g., Peltier, 2004),likely related to both the Laurentide and Cordilleran ice sheets.Where records extend beyond ~6 ka with index points and/or acombination of marine and terrestrial limiting points (#10 and#12), a decreasing rate through time is apparent due to thecombined effects of the exponential decrease in GIA and thedecreasing rate of eustatic sea-level rise after 7 ka (e.g., Lambeckand Chappell, 2001; Lambeck et al., 2014). These records providefurther evidence that the highest rates of RSL rise associated withforebulge collapse are distal to the edge of former ice sheets (e.g.,Peltier, 2001; Shennan and Horton, 2002; Davis et al., 2008;Engelhart and Horton, 2012).

6.2. Late Holocene sea-level history

Elastic dislocation models of the earthquake deformation cyclecommonly assume that present day GIA at the Cascadia subduc-tion zone is negligible (e.g., Wang et al., 2003, 2013; Wang, 2007).This assumption has been based on RSL data from isolation basinsin British Columbia spanning the time period from ~14 to ~4 ka(James et al., 2000). The data have been interpreted to demon-strate ~0.1 mm a�1 of ongoing GIA uplift (e.g., James et al., 2009;Roe et al., 2013), which is ascribed to a low viscosity mantlerelated to the subduction zone. Whilst the data employed in Jameset al. (2009) were spatially limited to southern Vancouver Island,the conclusion of minimal ongoing GIA has been extrapolatedbeyond the area of maximum extent of the Cordilleran ice sheet.The minimal ongoing GIA is problematic as viscosity structureobtained from formerly glaciated areas may only be indicative ofthe loaded region rather than a regional or global average (e.g.,Paulson et al., 2005).

During the late Holocene (last 4 ka), eustatic sea-level rise waszero (Peltier, 2002; Peltier et al., 2002; Bassett et al., 2005; Milneet al., 2005) or minimal (0.1e0.2 mm a�1, Lambeck, 2002;Lambeck et al., 2014) making GIA the predominant control on RSLchange (e.g., Milne et al., 2005; Church et al., 2008; Lambeck et al.,

2014). For this reason, we assume reconstructions of RSL during thelate Holocene are an approximation for rates of GIA-related land-level and geoid changes (e.g., Shennan and Horton, 2002; Engelhartet al., 2009; Engelhart and Horton, 2012).

Ongoing uplift (i.e. negative RSL rise) is documented on thewestern side of Vancouver Island (�0.7 ± 0.4 mm a�1), the north-western Georgia Strait (�0.4 ± 0.4 mm a�1) and in Queen CharlotteStrait (�0.7 ± 0.8 mm a�1; Fig. 5a,b). These results are similar tothose of James et al. (2005) who estimated late Holocene uplift of0.25 mm a�1 for eastern Vancouver Island. James et al. (2009)suggested a southwards decrease in late Holocene crustal upliftrates, with minimal rates of vertical motion in regions located nearthe periphery of the former ice sheet. However, our data show thatthere is no clear pattern in the east to west gradient of late Holo-cene rates, which may reflect complex deglacial histories or morecomplicated patterns of deformation than implied by the classicmodel of the earthquake cycle in the Puget Lowland/Georgia Straitregion (e.g., Wang, 2007).

Between 49� N and 47� N, index points show a change fromuplift to stability or low rates of subsidence (positive RSL rise,Fig. 5b). Rates vary from 0.0± 0.4 on the northernWashington coastto 0.7 ± 0.3 mm a�1 in the Puget Lowland and 0.7 ± 0.2 mm a�1 onthe southern Washington coast (Fig. 5a).

In regions beyond the maximum extent of the Cordilleran icesheet, estimates of late Holocene RSL rise increase with distancefrom the former ice sheet (Fig. 5a,b). RSL rise increases to 1 mm a�1

along the Oregon coast (1.1 ± 0.2 mm a�1 in the north and1.2 ± 0.3 mm a�1 in the central region). The highest rates of lateHolocene RSL rise are found south of 41�N, with the northern andcentral California coasts recording 1.5 ± 0.3 mm a�1. These resultsare in general agreement with average rates in Washington(0.7 ± 0.03 mm a�1) and Oregon (1.1 ± 0.07 mm a�1) proposed byLong and Shennan (1998). Long and Shennan (1998) suggested asystematic northesouth increase in GIA of 0.25 mm a�1 100 km�1.Analysis of our expanded database suggests that the northesouthincrease in GIA is likely lower, falling between 0.1 and 0.15 mm a�1

100 km�1.The ongoing subsidence along the coasts ofWashington, Oregon

and California, presumably due to GIA, is of similar magnitude tothat observed along the U.S. Atlantic coast (e.g., Engelhart et al.,2009; Engelhart and Horton, 2012) and in other regions (e.g.,Shennan and Horton, 2002; Vink et al., 2007) where proglacialforebulges have collapsed. GPS observations of vertical landmovement at Cascadia commonly assume that the majority ofvertical motion is due to deformation from the subduction zone(e.g., McCaffrey et al., 2013; Schmalzle et al., 2014). However, ouranalysis confirms that calculations of vertical land motion due tosubduction include significant components of GIA. To extract 20thcentury rates of sea-level rise from satellite altimeters and long-term tide-gauge records, corrections will need to be applied forvertical movements that are primarily a crustal response to GIA.

7. Conclusions

Along the Pacific coast of central North America, 600 indexpoints and 241 limiting points indicate two RSL patterns based ondifferent deglacial histories. Our compilation spans 16 ka to pre-sent, but, as with similar databases from the east coast of the UnitedStates and Europe, the majority of index points (95%) and limitingdata (54%) are younger than 7 ka. Sites formerly glaciated by theCordilleran Ice Sheet in British Columbia and Washington havenon-monotonic RSL curves, reflecting the dominant influence ofglacial isostatic adjustment and the interplay with changing ratesof global meltwater input. Where data are available, these eightregions demonstrate a fall in RSL immediately post-deglaciation

Fig. 5. (a) Estimated rates of late Holocene relative sea-level rise (mm a�1) with 2s errors for 12 locations (red dots) along the Pacific coast of North America. Inset plots areexamples of locations with sea-level index points plotted as calibrated age versus relative sea level, where 0 is present elevation. Solid red line is the linear regression with 2s errors(dashed red lines) for each site. (b) Late Holocene relative sea-level rise field generated by the empirical-Bayesian spatio-temporal statistical model for the entire study area. Thisfield is subsampled to obtain rates for our individual regions that are shown in (a). White diamonds represent all the sites in the database.

S.E. Engelhart et al. / Quaternary Science Reviews 113 (2015) 78e92 89

with a highstand in the mid and late Holocene. Sites at the pe-riphery or distant from the ice sheet show a monotonic form, withRSL slowing rising through time as a product of the decreasing rateof forebulge subsidence and the decrease in meltwater input at 7and 4 ka. Analysis of late Holocene RSL data demonstrates spatiallyvariable rates, from a fall of 0.7 ± 0.8 mm a�1 in British Columbia toa rise of 1.5 ± 0.3 mm a�1 in San Francisco Bay. We propose that thisvariability is primarily due to crustal response to the deglaciation ofthe Cordilleran and Laurentide Ice Sheets. Future GPS studies ofCascadia subduction zone deformation need to incorporate a GIAcorrection and must consider the late Holocene RSL historiesreconstructed from our index and limiting points in estimates ofongoing land-level changes.

Acknowledgments

This research was supported by NSF grants to SEE (EAR-1419844), BPH (EAR-1052848, EAR-1419824) and REK (ARC-1203415). MV contributes to the Labex OT-Med (n� ANR-11-LABX-0061) funded by the French government through the A*MIDEXproject (n� ANR-11-IDEX-0001-02). ARN is supported by theEarthquake Hazards Program of the U.S. Geological Survey (USGS).We thank Lee-Ann Bradley for constructing Fig. 1b and AndrewKemp for the USGS internal review. We are grateful to the associateeditor Jeff Freymueller and to the two anonymous reviewers fortheir comments that greatly improved the early version of thepaper. This paper is a contribution to IGCP Project 588 andPALSEA2.

Appendix A. Supplementary data

Supplementary data related to this article can be found at http://dx.doi.org/10.1016/j.quascirev.2014.12.001.

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