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

Quaternary Science Reviews xxx (2014) 1e15

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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, USA Aix-Marseille Universit
e, CEREGE CNRS-IRD UMR 34, Europole de l'Arbois BP 80, 13545 Aix-en-Provence Cedex 4, France c Sea Level Research, Department of Marine and Coastal Sciences, Rutgers University, New Brunswick, NJ 08901, USA d Institute of Earth, Ocean and Atmospheric Sciences, Rutgers University, New Brunswick, NJ 08901, USA e Division of Earth Sciences and Earth Observatory of Singapore, Nanyang Technological University, 639798, Singapore f Geologic Hazards Science Center, U.S. Geological Survey, Golden, CO 80401, USA g Department of Earth & Planetary Sciences and Rutgers Energy Institute, Rutgers University, Piscataway, NJ 08854, USA b

a r t i c l e i n f o

a b s t r a c t

Article history: Received 30 May 2014 Received in revised form 28 November 2014 Accepted 7 December 2014 Available online xxx

A database of published and new relative sea-level (RSL) data for the past 16 ka constrains the sea-level histories 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 geological and archaeological investigations yields 600 sea-level index points and 241 sea-level limiting points. We subdivided the database into 12 regions based on the availability of data, tectonic setting, and distance from the former Cordilleran ice sheet. Most index (95%) and limiting points (54%) are <7 ka; older data come mainly from British Columbia and San Francisco Bay. The stratigraphic position of points was used as a first-order assessment of compaction. Formerly glaciated areas show variable RSL change; where data 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 a decreasing rate through time due to the collapse of the peripheral forebulge and the reduction in meltwater input during deglaciation. Late Holocene RSL change varies spatially from falling at 0.7 ± 0.8 mm a
1 in southern British Columbia to rising at 1.5 ± 0.3 mm a
1 in California. The different sea-level histories are an ongoing isostatic response to deglaciation of the Cordilleran and Laurentide Ice Sheets. © 2014 Elsevier Ltd. All rights reserved.

Keywords: Sea-level database Cascadia subduction zone Pacific North America Glacial isostatic adjustment Holocene

1. Introduction Regional databases of relative sea level (RSL)dfor example, in Great 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 our understanding of the primary mechanisms of RSL change since the Last Glacial Maximum (~26 ka, e.g., Peltier et al., 2002). Regional databases also represent a long-term baseline against which to gauge changes in sea level during the 20th century (e.g., Mazzotti et al., 2008; Engelhart et al., 2009), forecasts for the 21st century (e.g., Horton et al., 2014), and the basis for identifying regional

* Corresponding author. Tel.: þ1 401 874 2187. E-mail address: [email protected] (S.E. Engelhart).

variations in RSL (e.g., Shennan and Horton, 2002). Further, deglacial RSL reconstructions are used to constrain geophysical models of glacial isostatic adjustment (GIA, e.g., Peltier et al., 2002; Mitrovica, 2003; Lambeck et al., 2004; Milne and Peros, 2013). Late-Holocene data are crucial to assess spatial variability of rates of ongoing GIA (e.g., Engelhart et al., 2009). Such information is particularly important to correct sea-level measurements obtained by instrumental methods (e.g., Church and White, 2011). Changes in RSL are the net effect of simultaneous contributions from eustatic, isostatic (glacio and hydro), tectonic and local factors, all of which have different response timescales. The relative importance of these factors varies in time and space along the central Pacific coast of North America. The greatest RSL change since the Last Glacial Maximum was caused by the melting of approximately 50 million km3 of ice in land-based ice sheets, raising RSL in regions distant from the major glaciation centers (far-

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Please cite this article in press as: Engelhart, S.E., et al., A sea-level database for the Pacific coast of central North America, Quaternary Science Reviews (2014), http://dx.doi.org/10.1016/j.quascirev.2014.12.001

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S.E. Engelhart et al. / Quaternary Science Reviews xxx (2014) 1e15

field sites) by ~120e130 m (e.g., Lambeck, 2002; Peltier and Fairbanks, 2006; Lambeck et al., 2014). This meltwater (or eustatic) contribution to far-field RSL rise during deglaciation averaged 12 mm a
1, although peak rates potentially exceeded 40 mm a
1 during “meltwater pulses” at 19 and 14.5 ka (e.g., Deschamps et al., 2012; Lambeck et al., 2014). Empirical and glacial isostatic adjustment (GIA) modeling studies suggest a significant reduction 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 ice sheets (near-field sites) as a consequence of glacial isostatic adjustment, or isostatic rebound of Earth's crust (e.g., Clague et al., 1982; Roe et al., 2013). Growth and thickening of an ice sheet results in subsidence of land beneath the ice mass, which is compensated for by an outward flow of mantle material that uplifts a peripheral bulge around the ice margin. When the ice sheet melts and loading is diminished, land beneath the melted ice is uplifted at rates which may locally reach 50e100 mm a
1 (e.g., Shaw et al., 2002). The peripheral forebulge subsides and moves progressively toward the center of the diminishing load as mantle material is redistributed (Peltier, 2004). Along parts of the Pacific coast of central North America (Fig. 1a), RSL histories are also of fundamental importance in recording vertical 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 cycles of strain accumulation and release on the megathrust fault where

the Juan de Fuca plate subducts beneath North America at the Cascadia subduction zone (e.g., Wells and Simpson, 2001; Wang et al., 2012; McCaffrey et al., 2013; Fig. 1b). Such histories of tectonic land-level change, in turn, have helped build more realistic models of plate-boundary and upper-plate deformation and more accurate assessments of the hazard posed by earthquakes on local and regional faults (e.g., Burgette et al., 2009; Wang et al., 2013; Nelson et al., 2014). As discussed for each region below, vertical rates of uplift or subsidencedintegrated over the centuries to millennia spanned by the data points of this compilationdare only high enough to influence rates of RSL change near Holocene faults or folds with high rates of deformation in the shallow crust (<25 km depth; e.g., James et al., 2009). Local but highly spatially variable factors may influence RSL history. These include changes in the tidal regime (e.g., Shennan et al., 2000a; Uehara et al., 2006; Hill et al., 2011; Hall et al., 2013; Horton et al., 2013a) due to changes in local geomorphology (e.g., Shennan et al., 2003) or to global changes in tidal amplitudes. The latter may be the result of changes in the availability 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 overlying sediment and land drainage (e.g., Kaye and Barghoorn, 1964; €rnqvist et al., 2008) can also produce significant errors in RSL To reconstructions. After summarizing previous Holocene sea-level studies along the Pacific coast of central North America, we explain how a

Fig. 1. (a) Approximate spatial extent of the Cordilleran and Laurentide ice sheet at ~ 21.5 ka redrawn from Clague and James (2002) and Dyke et al. (2002). b) Simplified tectonics of central western North America showing active faults and plate motions (based on Wells and Simpson, 2001; Kelsey et al., 2012; McCaffrey et al., 2013). Gray arrows show direction of modeled block movements from McCaffrey et al. (2013). Oblique subduction of the Juan de Fuca plate beneath North America along the Cascadia subduction zone (CSZ) at 35 mm a
1 causes rotation and northward movement of the Oregon Coast Range block against the buttress rocks of Vancouver Island (southward pointing short arrows) producing compression with active faulting in the Puget Lowland of northwest Washington. The San Andreas Fault (SAF) accommodates most of the motion between the Pacific plate and the Sierra Nevada block. Numbered rectangles show the location of sea-level data of this paper grouped into regions as explained in text. B&R e Basin and Range; NCS e Northern Cascadia; ORC e Oregon Coast Range; SNV e Sierra Nevada; and YAK e Yakima Fold Belt.

Please cite this article in press as: Engelhart, S.E., et al., A sea-level database for the Pacific coast of central North America, Quaternary Science Reviews (2014), http://dx.doi.org/10.1016/j.quascirev.2014.12.001

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consistent methodology for selecting sea-level index points improves on previous compilations, and outline our reasons for grouping sea-level data into 12 regions. We reevaluated the 342 sea-level data points compiled by Hutchinson (1992), added 499 new sea-level data points from more recent studies, for a total of 600 index points and additional 241 sea-level limiting points that provide maximum or minimum constraints on former sea levels. We then reconstruct and compare the RSL histories among different regions and briefly discuss some implications of the history of RSL changes along the coast of central North America. 2. Previous work The post-Last Glacial Maximum RSL history of southern British Columbia 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 and James, 2002; Hutchinson et al., 2004). These studies, particularly focused on the early to mid Holocene RSL history and on the magnitude of the sea-level lowstand, highlight the complex isostatic response of the area to downwasting and retreat of the Cordilleran Ice Sheet. The timing and magnitude of the sea-level lowstand, which have important implications for understanding different 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 coastal investigations with other goals, for example, studies of the earthquake and tsunami history of the Cascadia subduction zone, along the coasts of British Columbia, Washington, Oregon, and northern California (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 the Puget Lowland of northwest Washington supplied additional sealevel data (e.g., Bucknam et al., 1992; Sherrod et al., 2000; Kelsey et al., 2004, 2012). Paleo-tsunami investigations provided data points 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-level data have focused on shoreline development and sediment transport (e.g., Peterson et al., 1984, 2010; Williams and Roberts, 1989; Peterson and Phipps, 1992), geo-archaeology (e.g., Friele and Hutchinson, 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 and Shennan, 1998). Post-Last Glacial Maximum sea-level databases were recently published by Shugar et al. (2014) and McLaren et al. (2014) from southern Oregon to Alaska, and for the central Pacific coast of Canada, respectively. Our new database includes two additional regions (northern and central California) and additional data from Oregon and Washington (221 sea-level index points and limiting data). We also assess all index points more rigorously than Shugar et 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 and limiting points has been defined by the International Geoscience Programme (IGCP) projects 61, 200, 495 and 588 (e.g., Preuss, 1979; Tooley, 1982; van de Plassche, 1982; Shennan, 1987; Gehrels and Long, 2007; Horton et al., 2009; Switzer et al., 2012; Shennan et al., 2015). A sea-level index point is a discrete reconstruction of the unique position of RSL in time and space (van de Plassche et al.,

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2014) Where a suite of sea-level index points exists for a locality or region, they describe changes in RSL through time and estimate rates of change. A limiting point indicates formation in a freshwater or marine environment (Engelhart et al., 2011). Therefore, reconstructed RSL must fall below freshwater limiting dates and above marine limiting dates (Shennan and Horton, 2002). Although these data cannot be used to produce valid sea-level index points, they are extremely important in understanding and interpreting sealevel changes along the Pacific coast of central North America. We validated all index points and limiting points in our database by insuring that each sea-level indicator had (1) a location; (2) a calibrated age (in solar years); (3) an elevation at which the indicator was sampled, and (4) a known relationship between the indicator and a contemporaneous tidal level (termed the indicative meaning; van de Plassche, 1986; Shennan et al., 2015). Sea-level indicators are physical, biological or chemical features possessing a systematic and quantifiable relationship to elevation in the tidal frame (Shennan, 1986; van de Plassche, 1986). This methodology improves upon data from some previous compilations (e.g., Shugar et al., 2014) where only a generalized tidal environment (e.g., intertidal, subtidal, supratidal) was listed for many indicators and no indicator-specific elevational range or errors were estimated. For example, 342 of the data points in our database were extracted from the comprehensive catalogue of radiocarbon ages of Hutchinson (1992) based on data collected between Queen Charlotte Strait and Cape Mendocino. We re-evaluated the ages used by Hutchinson (1992) and added 500 additional data points from published and unpublished sources to our database. Many points compiled 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 from studies focused on dating the timing of great earthquakes and tsunamis along the Cascadia subduction zone. Almost all (>95%) such dated indicators yield maximum ages for the times of earthquakes and tsunamis, because the ages are on either (1) detrital materials (wood, herb parts, seeds) from the upper few centimeters of subsided wetland soils or from the tsunami or tidal deposits immediately above them, or plant stems or rhizomes rooted in the upper parts of the soils (e.g., Atwater and Yamaguchi, 1991; Nelson, 1992; Kemp et al., 2013). If, as is commonly assumed, coseismically subsided wetland soils recover much of their pre-earthquake elevation 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 earthquake provide accurate sea-level index points. Except in a few areas, such as near Cape Blanco (Kelsey, 1990) and Cape Mendocino (Merritts and Bull, 1989), studies of late Pleistocene marine terraces and rates of incision in the central Coast Range suggest variable but low rates of uplift along much of the central Pacific coast over tens of thousands 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 little affect on the rates of postglacial sea-level rise that we report. We excluded data points that may be significantly affected by non-elastic tectonic uplift or subsidence, such as from sites within a few kilometers of Holocene faults and folds with high rates of deformation in the Puget Lowland. For example, detailed RSL data astride 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 other cases, however, sea-level data from fault studies have probably not been deformed at rates high enough to invalidate them for our sealevel compilation. For example, Kelsey et al. (2004) compared rates of sea-level rise at two salt marshes on opposite sides of an obliqueslip Holocene fault on Whidbey Island. Sea-level index points

Please cite this article in press as: Engelhart, S.E., et al., A sea-level database for the Pacific coast of central North America, Quaternary Science Reviews (2014), http://dx.doi.org/10.1016/j.quascirev.2014.12.001

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younger than 2.5 ka on either side of the fault show a similar rate of rise and are used in our database. Similarly, in the Humboldt Bay area of northern California, Holocene slip rates range from ~6 mm a
1 on the Little Salmon fault and 5e7 mm a
1 on the Mad River fault (Clarke and Carver, 1992; McCrory, 2000; Hogarth et al., 2012) and large upper-plate earthquakes are documented in the area (Valentine et al., 2012). However, the sea-level index points near these faults that we used in our database (see section 5.11) did not show anomalous elevations, as did data points near the Seattle fault in the Puget Lowland. San Francisco Bay is the other of our regions where rates of uplift or subsidence on active Holocene faults may influence the elevation of sea-level data. Whereas Ryan et al. (2008) identified subsidence between 0.2 and 0.3 mm a
1 with maximal rates northwest of the Golden Gate area, the majority of the long-term slip on Golden Gate platform faults caused only minor regional uplift with a maximum rate of 0.025 mm a
1. Our San Francisco Bay sea-level index points come from areas 70 km to the east and 60 km to the south of the Golden Gate platform faults. Slip rates on faults in these areas are much lower than rates on the main faults of the San Andreas system and all faults are strike-slip (Dawson and Weldon, 2013). Thus, tectonic subsidence or uplift near our index points is probably less than 0.02 m ka
1. The final database includes 600 index points and 241 limiting points (Fig. 2a), 42 of which are unpublished. The age range of the data spans the last 16 ka, but the majority of the index points (95%) and limiting points (54%) are younger than 7 ka (Fig. 2b). Index points and limiting points older than 7 ka are generally found at sites in British Columbia, the Puget Lowland and San Francisco Bay. The complete database and the references to the original data are in Appendix 1 and Appendix 2, respectively. 3.1. Types of sea-level points Although a variety of types of geomorphic and stratigraphic evidence 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 database are from salt and freshwater wetlands or adjacent estuarine sediment. Following Engelhart and Horton (2012), three indicative meanings (Table 1) were utilized for sea-level indicators from saltmarsh sediment based on the zonation of vascular plant communities (e.g., Nelson and Kashima, 1993) and assemblages of microfossils (e.g., Jennings and Nelson, 1992; Hemphill-Haley, 1995; Hawkes et al., 2010), and to a lesser extent on d13C values of bulk sediment and plant macrofossils (e.g., Engelhart et al., 2013). The indicative meaning is composed of a reference water level (e.g., Mean High Water, MHW) that defines the relation of the indicator to a contemporaneous tide level and is the midpoint of the indicative 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 sediment is between Highest Astronomical Tide (HAT) and Mean Tide Level (MTL). For indicators deposited in a high salt-marsh environment, the indicative range is restricted to HAT to MHW. For those from low salt-marsh environments, we apply an indicative range of MHW to MTL. If an indicator formed in a freshwater or marine environment, it is classified as a sea-level limiting point. Although limiting points constrain the position of RSL less precisely than index points, they are important in reconstructing sea-level changes because RSL must fall below freshwater limiting points and above marine limiting points (e.g., Shennan and Horton, 2002). Freshwater limiting points usually formed at an elevation above HAT, but in some cases they may have formed at elevations within the

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

intertidal zone due to rising groundwater tables (e.g., Jelgersma, 1961; Shennan et al., 2000a; Engelhart and Horton, 2012). Therefore, a reference water level of MTL has been employed in this analysis. Marine limiting points formed below MTL. In the formerly glaciated regions of British Columbia characterized by tens of meters of glacial isostatic uplift, isolation basins may provide precise sea-level index points (e.g., James et al., 2009; Roe et al., 2013). Isolation basins are natural rock depressions that were, at different times in their history, isolated from, or connected to, the sea (e.g., Shennan et al., 2000b; Long et al., 2006; Horton et al., 2013b). When RSL exceeded the elevation of a basin sill, the basin was inundated by marine water. Conversely, when RSL was lower than the sill, the basin was isolated from the sea and freshwater 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 freshwater deposits. The age of basin isolation is estimated by dating the contact between marine and freshwater sediment. Based upon extensive work in Europe (e.g., Lloyd, 2000; Shennan et al., 2000b; Long et al., 2006), we conservatively estimate an indicative range for the transition from marine to freshwater sediment to be HAT to MTL. The height of RSL at the time of isolation is precisely constrained by the elevation (with respect to modern sea level) of the basin sill and not by the elevation of the contact between marine and freshwater sediment (Horton et al., 2013b).

Please cite this article in press as: Engelhart, S.E., et al., A sea-level database for the Pacific coast of central North America, Quaternary Science Reviews (2014), http://dx.doi.org/10.1016/j.quascirev.2014.12.001

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Table 1 Summary 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 Astronomical Tide; MHHW e Mean Higher High Water; MHW e Mean High Water; and MTL e Mean Tide Level. Sample type Index Points High marsh environment

Low marsh environment Undifferentiated salt marsh environment

Isolation Basin Limiting Points Marine limiting

Freshwater limiting

Evidence

Reference water level

Indicative range

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) Low marsh plant macrofossils low marsh plant macrofossils. Foraminiferal and diatoms assemblages dominated by low marsh taxa (e.g., Scott et al., 1996; Hawkes et al., 2010). Author listing of unnamed salt marsh plant macrofossils or identification only to genus level (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). Sediments recording the switch between freshwater and marine sediments that are supported by changing diatoms assemblages (e.g., James et al.., 2009; Roe et al., 2013).

(HATeMHW)/2

HATeMHW

(MHWeMTL)/2

MHWeMTL

(HATeMTL)/2

HATeMTL

(HATeMTL)/2

HATeMTL

MTL

Below MTL

MTL

Above MTL

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 clastic sediment may be supported by diatom assemblages dominated by marine taxa (Shennan et al., 1998). Isolation basin sediments with marine diatoms (e.g. Hutchinson et al., 2004) and marine shell assemblages (e.g. Hutchinson et al., 2004; Mosher and Hewitt, 2004). In-situ tree stumps (e.g., Atwater and Hemphill-Haley, 1997) Peat that does not meet the above requirements to be classified as an index point (e.g., Shennan and Horton, 2002). Peat with freshwater diatoms. Freshwater isolation basin sediments (lacustrine gyttja) dominated by freshwater diatoms and freshwater shells assemblages (e.g. Hutchinson et al., 2004).

We also evaluated index and limiting points based on stratigraphy described in archaeological investigations in British Columbia, Washington, Oregon, and California. Evidence of coastal habitation (e.g., midden deposits) from Washington, Oregon, and northern California indicate occupation of most sites within the last 2 ka (Woodward et al., 1990; Minor and Grant, 1996). The native peoples of southern British Columbia shared an economy heavily based on maritime, intertidal and riverine resources (Hutchinson and McMillan, 1997) and archaeological investigations in this region span the last 6000 years (Grier et al., 2009). Sites of former villages are marked by extensive midden deposits, composed largely 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 uncertain factors in relating midden deposits to a contemporary tidal level, archaeological data are interpreted as freshwater limiting points with a reference water level of MTL. 3.2. Elevation of former sea levels RSL for each dated indicator is estimated using the following equation (Shennan and Horton, 2002):

RSLi ¼ Ei
RWLi

(1)

where Ei and RWLi are, respectively, the elevation and the reference water level of indicator i, both expressed relative to the same datum, MTL in our analysis. The associated total vertical uncertainty is expressed by the following equation (Shennan and Horton, 2002):


1=2 E ¼ e21 þ e22 þ e23 þ e2n …

(2)

where e1, en represent the sources of error for each index point including the indicative range. The additional errors are those associated with calculating the indicator elevation. Such errors can be 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 from which the indicator was collected (e.g., salt marsh; Engelhart and Horton, 2012). Sea-level indicators without elevational data were rejected. We included an error of ±0.1 m to account for the stability of benchmarks (National Geodetic Survey classification, Horton et al., 2009). Indicator thickness is also incorporated into the error term. For some older bulk 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 to the angle of boreholes as a function of indicator depth, taken in this €rnqvist et al., 2008). study as ±1% (To We do not include an error term for potential paleotidal range change (e.g., Uehara et al., 2006; Hill et al., 2011; Horton et al., 2013a) or sediment compaction (e.g., Edwards, 2006; Horton and Shennan, 2009). Sediment compaction results in a lowering of the index point relative to the initial depositional elevation and will result in an overestimation of RSL rise. Such processes have been noted, for example, for indicators from buried former marshes at the Salmon River estuary (Nelson et al., 2004) in comparison to those from buried former marshes at other Oregon estuaries (Long and Shennan, 1998). We make a first order assessment of sediment compaction by subdividing index points into basal and intercalated categories (Shennan and Horton, 2002; Horton and Shennan, 2009). Basal index points are from sediment that directly overlies incompressible substrate (Jelgersma, 1961). Intercalated index points are derived from beds of low-density, organic-rich sediment within a sequence of higher density, clastic units (Shennan and Horton, 2002) and are, therefore, particularly susceptible to compaction (Jelgersma, 1961). Where stratigraphic information was unavailable for an index point we conservatively interpreted it as intercalated. Index points from isolation basins are compaction free as the height of the bedrock sill, and not the height of the sampled sediment, controls the position of former RSL. 3.3. Age of sea-level indicators All ages in our database are radiocarbon (14C) ages on organic material from salt and freshwater marshes or marine shells. Indicators were dated using different methods of radiocarbon

Please cite this article in press as: Engelhart, S.E., et al., A sea-level database for the Pacific coast of central North America, Quaternary Science Reviews (2014), http://dx.doi.org/10.1016/j.quascirev.2014.12.001

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analysis, including Gas Scintillation Counting (GSC), Liquid Scintillation 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 (Reimer et al., 2013) datasets for terrestrial and marine indicators, respectively. For shells we used carbon reservoir corrections from the Marine Reservoir Database (Reimer and Reimer, 2001) or from other published values for the region (Appendix 1). All index points are presented as calibrated years before present (ka), where year 0 is 1950 CE. Basal gyttja ages can be affected by the introduction of old carbon from underlying marine sediments (Hutchinson et al., 2004; James et al., 2009). A correction factor of
625 ± 60 years was applied to these ages (e.g., Hutchinson et al., 2004; James et al., 2009; Roe et al., 2013). Twentieth-century ages reported by the Geological Survey of Canada (GSC) radiocarbon lab are corrected to conventional ages following Stuiver and Polach (1977). A concern with radiocarbon ages analyzed prior to the mid-1970s is the correction used for isotopic fractionation (e.g., Hijma et al., 2015; €rnqvist et al., 2015). In the database, the majority of ages were To analyzed after 1980 and, therefore, most are not subject to this potential error. For the 37 ages that are affected, we followed the procedure of Hijma et al. (2015) to correct for isotopic fractionation. However, the general absence of C4 plant material within salt marshes on the Pacific coast of North America (e.g., Engelhart et al., 2013) minimizes this potential error in ages, as most dated material probably had values close to
25‰ relative to the VPDB standard. 3.4. Statistical framework to assess spatial variability of late Holocene relative sea-level change We assessed the spatial variability of GIA and tectonically induced land-level changes using an empirical-Bayesian spatiotemporal statistical model for late Holocene (<4 ka) sea-level change in which RSL at each site is assumed to change linearly over time at a rate m(x). The field m(x) has a Gaussian process prior (Rasmussen and Williams, 2006) with an exponential covariance function:

  
r ðx; x0 Þ mðxÞ  GP 0; s2s exp l

(3)

0

where s2s is the prior variance, rðx; x Þ the angular distance between x and x′, and l is a length scale. The index point observations yi are modeled (limiting data are not modeled):


 yi ¼ mðxi Þ tbi
t0 þ εti þ εyi þ wi þ yi gi þ y0 ðxi Þ:

(4)

The reference time t0 is set to 1970 CE. The term tbi is the mean estimated age of the observation, which has normally distributed error εti with variance specified for each point in the database. The noise term εyi 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 uncorrelated in space and time, with prior variance s2w ; this term allows the model to increase its noise estimate to improve the fit to the data. We allow for compaction in this model through the third error term, gi which is normally distributed with prior variance s2c . This term is multiplied by yi, so deeper samples, with more opportunity for compaction, have a larger error. The term y0(xi), which has a prior variance of s20 and is spatially uncorrelated, allows for offsets in the datums at different sites. We employ simulated annealing followed by local optimization to find the hyperparameters that maximize the likelihood of the model, 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 additional noise terms are essentially unneeded by the maximum-likelihood model, implying the vertical error specified in the database is adequate.) The rate estimates provided in this analysis are based on m(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 of northern and central California, Oregon, Washington, and southern British Columbia (Fig. 1b). Hutchinson (1992) subdivided his sealevel database into sub-regions according to latitude and longitude. Instead, we divided our database into sub-regions based on (1) regional tectonic setting, (2) distance from the former Cordilleran ice sheet, (3) proximity to other sea-level data points (commonly a function of local geomorphology, such as estuarine lowlands between coastal headlands), and (4) location relative to Holocene faults and folds. In simplified form, the regional-scale tectonics of westernmost central North America consists of shear interaction between the Pacific and North American plates south of the Mendocino Triple Junction, and oblique convergence between North America and the Juan de Fuca plate north of the Mendocino Triple Junction (Fig. 1b). South of the triple junction North America deforms through strikeslip faulting, largely along the San Andreas Fault system, whereas in the Cascadia forearc, north of the triple junction, plate subduction and the rotation of crustal blocks appear to accommodate the relative 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 clockwise rotation of the Cascadia forearc relative to North America (Wells and McCaffrey, 2013) and the most recent GPS data confirm that the rotation continues east of the forearc into the Basin and Range (McCaffrey et al., 2013). In northern Washington, motions of the Oregon coast block compress the Puget Lowland against the buttress of southwestern British Columbia. The current crustal shortening in the lowland of 4e5 mm a
1 is mostly accommodated by slip on eastewest, Holocene thrust faults a few tens of kilometers long that bound basins and uplifts (Fig. 1b, Blakely et al., 2002; Mazzotti et al., 2003; McCaffrey et al., 2013). In the San Francisco Bay region of central California, about 70e80% of the 45 mm a
1 of lateral slip between the Pacific and North America plates occurs along the complex faults of the San Andreas system (Grove et al., 2010; Tong et al., 2013). Vertical land-level changes (uplift and subsidence) along coasts of Cascadia and San Francisco Bay are largely the result of interseismic strain accumulation and release on faults of these regions during many earthquake cycles (Wang, 2007; Wang et al., 2012). The classic models of the earthquake deformation cycle invoke elastic deformation, all of which is assumed to be recoverable (Yeats et al., 1997; Scholtz, 2002). But at a particular site on a coastline, other factors such as RSL rise or fall, the amount of slip (proportional to earthquake magnitude) and how it was distributed on the fault plane (location of asperities, updip, downdip), the position of the site relative to the slip distribution, and viscoelastic effects over years to many decades following earthquakes, all modify these simple models of deformation (Wang, 2007; Wang et al., 2012; Nelson, 2013). At Cascadia, the megathrust is currently fully or close to fully locked, leading to margin-normal shortening and uplift of the forearc (Miller et al., 2001; Burgette et al., 2009). This leads to significant variability in ongoing crustal strain rates, especially in the northern portion of the subduction zone (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 beneath the wetlands of the San Francisco Bay region makes their postglacial vertical slip rates difficult to discern. In most cases, vertical rates are probably at least an order of magnitude lower than lateral rates, that is, <0.5 mm a
1 (e.g., Grove et al., 2010). Late Pleistocene uplift rates measured from an 80-ka terrace along a fault northwest of San Francisco are 0.2e1.0 mm a
1 (Grove et al., 2010), whereas Ryan et al. (2008) measured subsidence rate of 0.2e0.3 mm a
1 on a block bounding the San Andreas Fault system west of San Francisco Bay. Marsh elevations in south San Francisco Bay have kept pace with tens of centimeters of sediment compaction and subsidence due to groundwater withdrawal, and to a much lesser extent to sea-level rise and tectonic uplift or subsidence, over at least the past half century at rates as high as 40 mm a
1 (Patrick and DeLaune, 1990). The Cordilleran ice sheet formed over British Columbia, southern Yukon Territory and southern Alaska (Fig. 1a; Clague and James, 2002). Its glacial history, including the timing of the Last Glacial Maximum, differs from that of the Laurentide ice-sheet (Gregoire et al., 2012). At 21.4 ka, ice cover in North America was at or near its maximum everywhere except along the southern Cordilleran ice margin (Dyke, 2004). In particular, the southwestern portion of the Cordilleran ice-sheet continued to build to its maximum extent until ~17.2 ka (Clague and James, 2002) with expansion towards the coast along the Strait of Georgia, as expressed by the Puget and Juan de Fuca lobes (Clague, 1983; James et al., 2009; Fig. 1a). Deglaciation proceeded very rapidly along the southwestern Cordilleran ice margin; Juan de Fuca Strait and the Strait of Georgia were free of ice by ~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 northernmost region of the database. It is located between two plate triple junctions that mark the limits of the Explorer microplate (Audet et al., 2008; Fig. 1a). GPS-derived velocity fields indicate a significant deviation from the typical pattern of elastic strain rates resulting from locking on the subduction zone (McCaffrey et al., 2007). South of 50 N, GPS vertical velocities show an eastewest gradient with sites on the west coast indicating uplift at 1e4 mm a
1 and sites more inland indicating near-zero vertical motion or slight subsidence (Mazzotti et al., 2008). Therefore, we subdivided the database on the basis of these modern observations. On western Vancouver Island (#2), GPS vertical velocities show uplift rates ranging between ~2.6 and ~3.9 mm a
1. They decrease down 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 Washington coast (#6) and Puget Sound (#7) represent the southernmost limit of the former ice sheet (Clague, 1983; Clague and James, 2002; James et al., 2009). In this area, a permanent northesouth shortening is superimposed on the transient subduction signal (Hyndman et al., 2003; Mazzotti et al., 2003) and GPS vertical velocities show high variability. They range from ~3.9 mm a
1 of uplift on the northwestern Washington coast to zero vertical motion or slight subsidence in the Puget Lowland area (Mazzotti et al., 2008). The remaining regions (#8 to #12, Fig 1b) were not covered by the Cordilleran ice sheet during the Wisconsinan glaciation (Clague, 1983; Hutchinson et al., 2004). The southern Washington region (#8) marks the transition between the region of northesouth shortening and the region of northwards motion of the Oregon block. Here, GPS vertical velocities show the highest uplift with rates >2.7 mm a
1 (Mazzotti et al., 2008). The pattern, due to the locking of the subduction plate boundary along the southern

7

portion of the Cascadia margin, results in variability in uplift rates along the Oregon and northern California coast (Burgette et al., 2009; Schmalzle et al., 2014). Therefore, we subdivide the Oregon coast into two regions: the northern Oregon coast (#9), where GPS velocities indicate zero motion to slight subsidence, and the central Oregon coast (#10) where uplift rates of ~1.1 mm a
1 were recorded (Mazzotti et al., 2008). The northern California region (#11) is located roughly 100 km north of the Cape Mendocino triple junction, which marks the southern limit of the Cascadia subduction zone. Deformation in this region is dominated by GordaePacificeNorth America oblique plate convergence, with internal deformation of the Gorda plate, and lateral deformation along the Mendocino transform fault (Williams et al., 2006). Central California (#12) is the southernmost region of the database and is influenced primarily by strike-slip deformation along the San Andreas Fault system and related faulted systems to the south and east. 5. Sub-region sea-level histories 5.1. Queen Charlotte Strait (#1) The RSL history is composed of 4 index points and 13 limiting points (Fig. 3, #1) from the isolation basins of Seymour and Belize Inlets and offshore cores from Cook Bank (place names in Appendix 1). At ~13.0 ka a single index point places RSL at 1.3 ± 0.9 m. At ~12 ka RSL dropped dramatically to ~
95 m. There is an absence of index points during the early Holocene, although freshwater limiting points suggest RSL was below 3.6 m from ~12.0 to ~8.3 ka. At ~8.0 ka, one index point places RSL above present at ~1.3 m. The late Holocene 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 30 limiting points (Fig. 3, #2). Salt-marsh sediment provided index points from Tofino, Hesquiat Peninsula, Barkley Sound, Nootka Sound, Esperanza Inlet and Alberni Inlet, and isolation basins yielded index points and limiting dates from Effingham Inlet and Meares Island (place names in Appendix 1). Data from Lateglacial isolation basins record a very rapid RSL fall. The oldest marine limiting point indicates RSL was above 25 m at ~15.5 ka. Two limiting points constrain RSL to ~4 m MTL at ~14.0 ka, whereas index points indicate a sea level of ~
46.9 m between ~13.0 and ~12.3 ka. Freshwater limiting points constrain RSL to below
9 m at ~9.0 ka. Younger intercalated index points and freshwater limiting points indicate RSL was 2.5 ± 1.1 m MTL at ~6.8 ka, suggesting a rapid rise to a highstand during the early Holocene. Further intercalated 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, one intercalated index point and two limiting points constrain RSL to 2 ± 1.1 m MTL. Younger index points indicate RSL remained above present MTL until ~0.7 ka. 5.3. Northwestern Georgia Strait (#3) The RSL history of eastern Vancouver Island is constrained by 12 index points and 51 limiting points (Fig. 3, #3). Index points were from isolation basins on Lasqueti, Quadra and Cortes Islands and salt marshes near Qualicum and Courtenay (eastern Vancouver Island; place names in Appendix 1). Limiting samples were from archaeological excavations near Comox and Nanaimo, from isolation basins on Lasqueti, Quadra and Cortes Islands, and from the coast of eastern Vancouver Island. Marine and freshwater limiting points constrain RSL to ~150.0 m at ~13.9 ka. Multiple terrestrial

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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 (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 and age errors. BC e British Columbia, Canada; WA e Washington, USA; CSZ e Cascadia Subduction Zone; Qcs e Queen Charlotte Strait; Gs e Georgia Strait; Jfs e Juan de Fuca Strait; Ps e Puget Sound.

and marine limiting points then show a rapid sea-level fall to an index point with a RSL of
0.9 ± 1.1 m at ~12.0 ka. Marine limiting points 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 to between
20 and 0 m. At ~8.1 ka, one marine limiting point suggests RSL was above present. Intercalated index points indicate RSL was at 1.0 m ± 1.1 at ~6.5 ka and 2.5 ± 1.1 m at ~3.2 ka. Younger index 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 23 index points and 42 limiting points (Fig. 3, #4). Archaeological excavations and salt and freshwater marshes from the Fraser River delta, Howe Sound, Padilla Bay and Bellingham Bay provided the data (place names in Appendix 1). Additional limiting points were from the glaciomarine deposits in the Fraser River Valley and from

freshwater marshes and archaeological excavations in the Gulf Islands and Skagit Bay. Marine and terrestrial limiting points indicate a rapid fall in RSL at the transition from the Lateglacial period to the early Holocene. A single marine limiting point suggests that RSL was above 154 m at ~14.6 ka. At ~13.3 ka, RSL was between 66 and 58 m, and between ~12.0 and ~10.0 ka limiting points constrain RSL to between 12 m and
33 m. At ~9.2 ka, RSL was below
11 m. 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 the late Holocene. Intercalated index points of late Holocene age show scatter, likely illustrating the influence of compaction. The youngest points 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 10 index points and 50 limiting points (Fig. 3, #5). Isolation basins

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(near Victoria and along eastern Juan de Fuca Strait), salt and freshwater marshes and archaeological excavations (Montague Harbor, 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 ± 1.3 m MTL at ~13.6 ka. In the early to mid Holocene (~12.0 to ~6.5 ka), RSL was between
42.5 and
1.4 m, and during the mid to late Holocene (~6.0 to ~3.3 ka), RSL was between
3.2 and 0 m. Two intercalated index points place RSL 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 14 index points and 3 limiting points (Fig. 3, #6) and is restricted to the last ~2.5 ka. Data came from the salt and freshwater marshes of the Waatch River, Crescent Bay and Neah Bay (place names in Appendix 1). The oldest index point at ~2.4 ka shows that RSL was at
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 m MTL) at ~0.6 ka. 5.7. Puget Sound (#7) RSL history for the Puget Lowland region is restricted to the Lateglacial period and to the mid- and late Holocene (Fig. 3, #7). The database includes 35 index points and 9 limiting points from the marshes of Whidbey Island, Hood Canal, Quilcene Bay, Discovery Bay, the Nisqually delta, Cultus Bay, Bainbridge Island, and archaeological excavations east of Seattle (place names in Appendix 1). 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 to present RSL rose more slowly: intercalated index points indicate that 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 the early Holocene to the present. The data consist of 128 index points and 4 limiting points from salt marshes and estuarine sediments of Willapa Bay, Grays Harbor and the Columbia River (Fig. 4, #8). The oldest index point places RSL at
53 ± 1.3 m at ~10.0 ka. Then, RSL rapidly 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 intercalated index points, likely reflecting the influence of compaction. In the mid Holocene, RSL rose from
4.9 ± 0.4 at ~5.1 ka to
4.5 ± 1.0 m at ~4.6 ka. RSL in the late Holocene rose from
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 and freshwater marshes, estuarine sediments, and archaeological investigations of Tillamook Bay, Clatsop Plain, Salmon River estuary, Netarts Bay and Nehalem River estuary (Fig. 4, #9; place names in Appendix 1). In the early to mid-Holocene marine limiting data indicate RSL was above
28.7 m MTL at ~9.3 ka and above
19.3 m MTL at ~8.8 ka. Two intercalated index points suggest RSL was at
11.60 ± 1.0 m MTL at ~6.5 ka and at
10.6 ± 1 m MTL 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

9

index points. However, one basal index point suggests RSL was at
2 ± 1 m MTL at ~2.0 ka. 5.10. Central Oregon coast (#10) The Holocene RSL history of the central Oregon coast is constrained by 110 index points and 20 limiting points from the salt and freshwater marshes, and estuarine sediments of the Alsea River estuary, Yaquina Bay, Coos Bay, Coquille River, Siuslaw River, and archaeological investigations at Tahkenitch Lake (Fig. 4, #10; place names in Appendix 1). The oldest freshwater limiting point places RSL below
54.5 m MTL at ~11.5 ka. Marine limiting points suggest RSL was above
30.0 m MTL at ~9.5 ka,
18.9 m MTL at ~8.3 ka, and
12.6 m MTL at ~7.4 ka. Freshwater and marine limiting points constrain RSL to between
7 m and
10 m MTL at ~6.5 ka. One basal index point suggests that RSL was
7.9 ± 0.5 m MTL at ~6.6 ka, and another that it rose to
5.3 m ± 1.1 m MTL at ~4.6 ka. A third shows 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 late Holocene. The database is composed of 45 index points and 6 limiting point from the salt and freshwater marshes of Humboldt Bay and near Crescent City (Fig. 4, #11). The oldest basal index point at ~4.0 ka places RSL at
6.3 ± 1.0 m MTL, and anther show that it rose to
6.1 ± 0.1 m MTL at ~3.5 ka. There is a scatter in the index points between ~4.0 and ~2.0 ka. Most of the scatter is likely related to compaction rather than to uplift or subsidence adjacent to the Little Salmon and Mad River fault zones. Scatter in the index points decreases in the last ~2.0 ka. RSL was at
2.0 ± 0.1 m MTL at ~1.5 ka and 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 is composed of 70 index points from the salt marshes of the Sacramento River Delta and southern San Francisco Bay (Fig. 4, #12). RSL was
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 was at
9.8 ± 1.1 m MTL at ~7.8 ka and at
8.2 m ± 0.6 m at ~6.5 ka. Late Holocene 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 central North America Analysis of the RSL histories from 16 ka to present highlights significant differences between regions formerly covered by the Cordilleran ice-sheet and those located south of the ice margin. For the regions that were formerly covered by the Cordilleran ice sheet (Fig. 3, regions #1 to #7), RSL histories are strongly controlled by the ice unloading history, resulting in variable RSL histories due to the complex interplay of eustatic sea-level change and isostatic adjustment of the crust. This is typically expressed as a RSL fall following deglaciation due to isostatic uplift, followed by slow rates of RSL rise during the mid to late Holocene. Regions with records extending beyond the Holocene (Fig. 3, #1, #2, #3, #4; #5, #7) show the effects of GIA due to Cordilleran ice sheet retreat (e.g., Clague, 1980; Dethier et al., 1995; Clague and James, 2002) with RSL as much as 154 m above present at 14 ka. Significant differences in RSL (>45 m) exist at ~12 ka between Queen Charlotte Strait (#1) and western Vancouver Island (#2), which likely reflect the effects of

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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 and age errors. BC e British Columbia, Canada; WA e Washington; 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.

differing ice load history. Clague and James (2002) show that western Vancouver Island (#2) sites were deglaciated between ~18 and ~14 ka, earlier than sites in Queen Charlotte Strait (#1) that were still ice covered at ~14 ka. Both in northwestern Georgia Strait (#3) and southern Vancouver Island (#5) RSL falls below present between ~14 and ~12 ka before rising until 7 ka due to decreasing isostatic uplift and continuing high rates of eustatic sea-level rise (Lambeck et al., 2014). Our database provides insight into the variability in the reported depth of the sea-level lowstand across the region following deglaciation in British Columbia (James et al., 2009). In southern Vancouver Island (#5), James et al. (2009) placed the lowstand at
30 ± 5 m at ~11 ka. This depth is ~20 m shallower than the lowstand depth previously proposed (Linden and Schurer, 1988; Mosher and Hewitt, 2004). The variability in lowstand elevation is due to different interpretations of marine terraces identified by seismic data that suggest differences in ice loading, and hence GIA history. Along the western coast of Vancouver Island (#2),

isolation basin-derived index points from Effingham Inlet suggest a lowstand of ~
47 m at ~13 ka. This is a deeper lowstand than found in eastern Vancouver Island (#3) and southeastern Georgia Strait (#4), where marine limiting points constrain the lowstand to above ~
11 and ~
33 m from ~12 to ~10 ka. James et al. (2009) demonstrated that these differences are expected due to differential isostatic effects. Less definite conclusions can be drawn about the characteristics of the lowstand along southern Vancouver Island (#5). Marine limiting points indicate RSL was above
42 m at ~14.5 ka and above
32 m at ~9 ka, but these data do not preclude a higher or lower RSL between these ages. The lack of early Holocene data precludes identifying the elevations of lowstands along the northern Washington and Puget Lowland coasts (#6, #7). Four of the formerly glaciated regions had mid-Holocene highstands, which varied in elevation and temporally. In Queen Charlotte Strait (#1), RSL rose to ~1.3 m at ~8.0 ka. At western Vancouver Island (#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 an intercalated index point on Vargas Island. At northwestern Georgia Strait (#3), marine and terrestrial limiting data suggest RSL was above present between ~9 and ~8 ka, and an intercalated index point 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 of global meltwater input and the locally variable GIA. During the early Holocene, meltwater input outpaced the exponentially €rnqvist and Hijma, 2012). declining isostatic uplift rate (e.g., To However, with the decrease in meltwater input starting at ~8 ka (Lambeck and Chappell, 2001; Lambeck et al., 2014) due to the final melting of the Laurentide Ice Sheet (e.g., Carlson et al., 2008), isostasy became the dominant control on RSL history. This resulted first in a stable highstand above present sea level before RSL fell with a further decrease in meltwater input after ~4 ka (e.g., Peltier and Tushingham, 1991; Fleming et al., 1998; Milne et al., 2005; Engelhart et al., 2009). Similar RSL histories have been described from regions towards the periphery of ice sheet margins in northeastern Scotland (e.g., Shennan et al., 2002), but histories differ at sites in similar ice margin proximal settings in Maine (e.g., Engelhart and Horton, 2012) where no Holocene RSL highstand is recorded. The absence of a highstand with RSL rising continuously since ~6 ka at sites along southeastern Georgia Strait (#4) most likely represents persistent GIA-related subsidence during the mid and late Holocene. A lack of index points in southern Vancouver Island (#5) prevents a more detailed reconstruction of midHolocene RSL history. Sites from south of the Cordilleran ice margin (Fig. 4, #8 to #12) demonstrate a continuous rise of RSL throughout the Holocene due 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 a combination of marine and terrestrial limiting points (#10 and #12), a decreasing rate through time is apparent due to the combined effects of the exponential decrease in GIA and the decreasing rate of eustatic sea-level rise after 7 ka (e.g., Lambeck and Chappell, 2001; Lambeck et al., 2014). These records provide further evidence that the highest rates of RSL rise associated with forebulge 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 cycle commonly assume that present day GIA at the Cascadia subduction zone is negligible (e.g., Wang et al., 2003, 2013; Wang, 2007). This assumption has been based on RSL data from isolation basins in British Columbia spanning the time period from ~14 to ~4 ka (James et al., 2000). The data have been interpreted to demonstrate ~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 mantle related to the subduction zone. Whilst the data employed in James et al. (2009) were spatially limited to southern Vancouver Island, the conclusion of minimal ongoing GIA has been extrapolated beyond the area of maximum extent of the Cordilleran ice sheet. The minimal ongoing GIA is problematic as viscosity structure obtained from formerly glaciated areas may only be indicative of the 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 was zero (Peltier, 2002; Peltier et al., 2002; Bassett et al., 2005; Milne et al., 2005) or minimal (0.1e0.2 mm a
1, Lambeck, 2002; Lambeck et al., 2014) making GIA the predominant control on RSL change (e.g., Milne et al., 2005; Church et al., 2008; Lambeck et al.,

11

2014). For this reason, we assume reconstructions of RSL during the late Holocene are an approximation for rates of GIA-related landlevel and geoid changes (e.g., Shennan and Horton, 2002; Engelhart et al., 2009; Engelhart and Horton, 2012). Ongoing uplift (i.e. negative RSL rise) is documented on the western side of Vancouver Island (
0.7 ± 0.4 mm a
1), the northwestern Georgia Strait (
0.4 ± 0.4 mm a
1) and in Queen Charlotte Strait (
0.7 ± 0.8 mm a
1; Fig. 5a,b). These results are similar to those of James et al. (2005) who estimated late Holocene uplift of 0.25 mm a
1 for eastern Vancouver Island. James et al. (2009) suggested a southwards decrease in late Holocene crustal uplift rates, with minimal rates of vertical motion in regions located near the periphery of the former ice sheet. However, our data show that there is no clear pattern in the east to west gradient of late Holocene rates, which may reflect complex deglacial histories or more complicated patterns of deformation than implied by the classic model of the earthquake cycle in the Puget Lowland/Georgia Strait region (e.g., Wang, 2007). Between 49 N and 47 N, index points show a change from uplift to stability or low rates of subsidence (positive RSL rise, Fig. 5b). Rates vary from 0.0 ± 0.4 on the northern Washington coast to 0.7 ± 0.3 mm a
1 in the Puget Lowland and 0.7 ± 0.2 mm a
1 on the southern Washington coast (Fig. 5a). In regions beyond the maximum extent of the Cordilleran ice sheet, estimates of late Holocene RSL rise increase with distance from 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 and 1.2 ± 0.3 mm a
1 in the central region). The highest rates of late Holocene RSL rise are found south of 41 N, with the northern and central California coasts recording 1.5 ± 0.3 mm a
1. These results are 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 by Long and Shennan (1998). Long and Shennan (1998) suggested a systematic northesouth increase in GIA of 0.25 mm a
1 100 km
1. Analysis of our expanded database suggests that the northesouth increase in GIA is likely lower, falling between 0.1 and 0.15 mm a
1 100 km
1. The ongoing subsidence along the coasts of Washington, Oregon and California, presumably due to GIA, is of similar magnitude to that 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 proglacial forebulges have collapsed. GPS observations of vertical land movement at Cascadia commonly assume that the majority of vertical motion is due to deformation from the subduction zone (e.g., McCaffrey et al., 2013; Schmalzle et al., 2014). However, our analysis confirms that calculations of vertical land motion due to subduction include significant components of GIA. To extract 20th century rates of sea-level rise from satellite altimeters and longterm tide-gauge records, corrections will need to be applied for vertical movements that are primarily a crustal response to GIA. 7. Conclusions Along the Pacific coast of central North America, 600 index points and 241 limiting points indicate two RSL patterns based on different deglacial histories. Our compilation spans 16 ka to present, but, as with similar databases from the east coast of the United States and Europe, the majority of index points (95%) and limiting data (54%) are younger than 7 ka. Sites formerly glaciated by the Cordilleran Ice Sheet in British Columbia and Washington have non-monotonic RSL curves, reflecting the dominant influence of glacial isostatic adjustment and the interplay with changing rates of global meltwater input. Where data are available, these eight regions demonstrate a fall in RSL immediately post-deglaciation

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S.E. Engelhart et al. / Quaternary Science Reviews xxx (2014) 1e15

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 are examples 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. This field is subsampled to obtain rates for our individual regions that are shown in (a). White diamonds represent all the sites in the database.

with a highstand in the mid and late Holocene. Sites at the periphery or distant from the ice sheet show a monotonic form, with RSL slowing rising through time as a product of the decreasing rate of forebulge subsidence and the decrease in meltwater input at 7 and 4 ka. Analysis of late Holocene RSL data demonstrates spatially variable rates, from a fall of 0.7 ± 0.8 mm a
1 in British Columbia to a rise of 1.5 ± 0.3 mm a
1 in San Francisco Bay. We propose that this variability is primarily due to crustal response to the deglaciation of the Cordilleran and Laurentide Ice Sheets. Future GPS studies of Cascadia subduction zone deformation need to incorporate a GIA correction and must consider the late Holocene RSL histories reconstructed from our index and limiting points in estimates of ongoing land-level changes. Acknowledgments This research was supported by NSF grants to SEE (EAR1419844), BPH (EAR-1052848, EAR-1419824) and REK (ARC1203415). MV contributes to the Labex OT-Med (n ANR-11-LABX0061) funded by the French government through the A*MIDEX project (n ANR-11-IDEX-0001-02). ARN is supported by the Earthquake Hazards Program of the U.S. Geological Survey (USGS). We thank Lee-Ann Bradley for constructing Fig. 1b and Andrew Kemp for the USGS internal review. We are grateful to the associate editor Jeff Freymueller and to the two anonymous reviewers for their comments that greatly improved the early version of the paper. This paper is a contribution to IGCP Project 588 and PALSEA2.

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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