Holocene coastal dune fields used as indicators of net littoral transport: West Coast, USA

Geomorphology 116 (2010) 115–134

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Holocene coastal dune fields used as indicators of net littoral transport: West Coast, USA Curt D. Peterson a,⁎, Errol Stock b, Roger Hart c, David Percy a, Steve W. Hostetler d, Jeffrey R. Knott e a

Portland State University, Portland, Oregon, 97207-0751, USA Griffith University, Brisbane, Queensland, 4111, Australia Oregon State University, Newport, Oregon, 97365, USA d U.S. Geological Survey, Oregon State University, Corvallis, Oregon, 97331, USA e California State University Fullerton, Fullerton, California, 93454, USA b c

a r t i c l e

i n f o

Article history: Received 24 September 2008 Received in revised form 12 October 2009 Accepted 18 October 2009 Available online 25 October 2009 Keywords: Coastal dunes Littoral drift West Coast USA Paleoclimate model Shoreline angle Wind-Wave forcing

a b s t r a c t Between Point Grenville, Washington, and Point Conception, California (1500 km distance) 21 dune fields record longshore transport in 20 littoral cells during the late Holocene. The direction of predominant littoral transport is established by relative positions of dune fields (north, central, or south) in 17 representative littoral cells. Dune field position is north of cell midpoints in northernmost Oregon and Washington, but is south of cell midpoints in southern Oregon and California. Downdrift sand trapping occurs at significant changes in shoreline angle and/or at bounding headlands that project at least 2.5 km seaward from the general coastal trend. Sand bypassing occurs around small headlands of less than 0.5 km in projection distance. A northward shift of the winter low-pressure center in the northeast Pacific Ocean is modeled from 11 ka to 0 ka. Nearshore current forcing in southern Oregon and northern California switched from northward in earliest Holocene time to southward in late Holocene time. The late Holocene (5–0 ka) is generally characterized by net northward littoral drift in northernmost Oregon and Washington and by net southward littoral drift in southernmost Oregon and California. A regional divergence of net transport direction in central Oregon, i.e. no net drift, is consistent with modeled wind and wave forcing at the present time (0 ka). © 2009 Elsevier B.V. All rights reserved.

1. Introduction The direction(s) of net littoral drift and resulting littoral sand transport to beaches of the US West Coast are addressed in this paper by using the asymmetries of coastal dune field distributions in representative littoral cells. The alongshore distributions of coastal dune fields are used to constrain long-term net transport in littoral cells from Washington, Oregon, and California, totaling about 1500 km in length (Fig. 1). Nearly 50 years ago Cooper (1958, 1967) mapped dune fields along the study area, but the dune field positions were not related to littoral sand sources, nearshore wave fields, or longshore currents. Several recent studies from other wave-dominated coastlines have used Holocene development of coastal dune fields to identify sources of littoral sediment supply and downdrift littoral transport in Brazil (Barbosa and Dominguez, 2004) and southern Africa (Bluck et al., 2007). The Holocene dune fields in the US West Coast study area (dated 7– 0 ka) generally formed during marine high-stand conditions (Orme, 1990; Weidemann, 1990; Jungner et al., 2001; Erlandson et al., 2005).

In this study we investigate the use of late Holocene dune fields as proxies to establish net littoral transport of beach sand in the US West Coast. The positions of 21 dune fields in 20 littoral cells from Washington to California (Fig. 1) are presented to demonstrate the net transport direction in each littoral cell. The presence or absence of dune fields is used as a proxy for sand residence time on beaches (Psuty, 1988; Hesp, 1999). Sand residence time, and corresponding coastal dune sand accumulation, should increase where beach sand is trapped against downdrift changes in shoreline angle and/or bounding headlands. The relative positions of dune fields in the study area littoral cells are compared to paleo-wind and wave forcing modeled at 11, 6 and 0 ka (Patrick and Hostetler, 2004). A general agreement between dune field positions within littoral cells, and the predicted direction of winter wave attack supports the use of dune field development in establishing dominant directions of longshore transport in littoral cells in this study area. The topics raised here should have direct relevance to other coastlines with sandy beaches and onshore winds. 2. Origins of coastal dune fields and study area setting

⁎ Corresponding author. Geology Dept. Portland State University, 1721 SW Broadway, Portland, Oregon, 97207-0751, USA. Tel.: +1 503 725 3375; fax: +1 503 725 3025. E-mail addresses: [email protected] (C.D. Peterson), [email protected] (E. Stock), [email protected] (R. Hart), [email protected] (D. Percy), [email protected] (S.W. Hostetler), [email protected] (J.R. Knott). 0169-555X/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2009.10.013

Shoreward movement of shelf sand by asymmetric wave transport occurred during the Holocene transgression on many coastlines that are characterized by broad continental shelves and high wave energy, such as in southeastern Australia (Thom et al., 1981) and south Africa

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Fig. 1. Map of the US West Coast study area extending from Point Grenville, Washington, to Point Arguello, California. The study area trends north–south (northern half) to north northwest–south southwest (southern half). Coordinates are in degrees Lat and Long, and in UTM-northing (N) and easting (E) at 100 km intervals. Selected headlands and geographic points are named (shown west of the coastline). Letters (A–U) designate shoreline map sections that are shown in Fig. 3. The largest rivers, which transport 1–5 × 106mt year−1 of bedload, are the Columbia, Klamath, Eel, Rogue, and Umpqua Rivers, in decreasing order of sediment supply. Accreted barriers and beach plains occur in the large Columbia River littoral cell (CRLC). A late Pleistocene shelf depocenter at Heceata Banks (open triangle) supplied sand to large dune fields in south-central Oregon.

(Illenberger and Verhagen, 1990). The long-term forcing of Holocene coastal dune development in the high wave energy coast of eastern Australia is directly linked to cross-shore sand supply that occurred during the Holocene transgression (Pye and Bowman, 1984). Initial shoreward sand transport in high-energy coastlines occurred soon after the decline in the rate of eustatic, sea-level rise at about 9-8 ka. The same relation has been shown for the largest dune fields in the US West Coast study area. Pleistocene sand depocenters on the continental shelf controlled the distribution of large, upland dune sheets in south-central Oregon (Peterson et al., 2007) (Fig. 1). The Oregon depocenter is located south of a major bight in the shelf that dammed northward littoral transport during sea-level low-stands. Eolian dune sands crossed the exposed low-stand shelf to reach the coastal foothills during late Pleistocene times. Beaches and smaller dune fields that are located on either side of the large shelf depocenter in central Oregon were also supplied by shoreward sand transport, as shown by

shelf heavy-mineral tracers (Clemens and Komar, 1988). During late Holocene times the nearshore currents redistributed available littoral sand to fill embayments, build spits, and locally feed coastal dune systems (Orme, 2002). Coastal dune fields are known to form downdrift from river mouths in Brazil (Barbosa and Dominguez, 2004) and south Africa (Bluck et al., 2007). Recent work on barrier spits and beach plains of the Columbia River littoral cell in the northernmost part of the US West Coast study area (Fig. 1) has confirmed their origins from downdrift supply of Columbia River mouth sand (Woxell, 1998; Herb 2000). The narrow beaches and smaller dune fields in southern Oregon and California are probably supplied by mixtures of both river sediment and remobilized shelf sand. Their relative proportions have yet to be evaluated on a regional basis. Of importance to this study, the positions of the smaller dune fields are not tied to river mouths (Cooper, 1958; 1967; Peterson et al., 1991; Orme, 1992; Dingler and Clifton, 1994). Wave energy is

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sufficient to disperse river sand supply away from the smaller river mouth sources in the sand-limited beaches of the study area. Though not directly addressed in this study, coastal dune development is typically episodic. Episodic Holocene dune development has been related to relative sea-level change in Brazil (Martin et al., 1998), and climatic changes in rainfall, wind power, and/or vegetation in northern Australia (Lees et al., 1995) and Israel (Gvirtzman et al., 1988). In this study we utilize long-term development of coastal dune fields (5–0 ka) to characterize longshore transport during the late Holocene. Efforts are underway to examine shorter-term processes of foredune development in the northern part of the study area (Ruggiero et al., 2005). For this paper we neglect the short-term processes of episodic and/or reversible sand transport to the beaches and foredunes that yield the larger dune fields over millennial timescales. The application of a broadly distributed proxy for discriminating net littoral transport is necessary to resolve conflicting reports of littoral drift in the US West Coast study area. In southernmost California (Fig. 1), Herron (1960) established southeastward net littoral transport from historic dredging records. A lack of consistent spit orientation in small bays of northern California and Oregon (Dingler and Clifton, 1994) has contributed to uncertainty of net littoral transport in that region. Based on symmetrical sand accumulation around several coastal jetties Komar et al. (1976) report zero net littoral drift in Oregon. Northward littoral transport is proposed for prograded barriers and beach plains located north of the Columbia River mouth in Washington (Ballard, 1964; Swartz, et al., 1985). Two types of Holocene dune fields are used to establish net littoral transport directions in the US West Coast region (Fig. 1). They include 1) upland dune fields that extend inland over marine terraces, flood plains, or Pleistocene dune sheets (Cooper, 1958, 1967; Hunter et al., 1983; Beckstrand, 2001), and 2) lowland dune fields of the Columbia River littoral cell that are characterized by abandoned foredune ridges in prograded beach plains (Cooper, 1958; Rankin, 1983; Woxell, 1998). The inland migration of upland dunes could be influenced by several factors including onshore wind conditions, topography, and vegetation. Those factors are not addressed here as only the shoreline positions of the dune fields are analyzed in this study (see Section 3). Significant-wave-heights in the study area currently reach 7–13 m during winter storms and 1–2 m during fair-weather summer months (Seymour, 1996; Tillotsen and Komar, 1997). Deepwater waves arrive from southwest or west during winter storms and northwest during fairweather conditions. Wave climate varies over inter-decadal periods, e.g., El Nino cycles, and possibly over multi-decadal periods (Peterson, et al., 1985, 1990; Inman and Jenkins, 1997; Allan and Komar, 2006). Deepwater wave directions in the eastern Pacific Ocean margin are expected to vary with latitude relative to low-pressure centers that generally travel from west to east. Regional sea-level pressure (SLP) is typically used to hindcast directional wave climate and shelf currents where instrumental data are not available. Independent measures of long-term net littoral forcing are needed to validate both the hindcast modeling and the very short-term trends analyzed from buoy and current meter data. Regional studies of heavy minerals from beach and river sand have been conducted in the study region to help detect directions of littoral transport. Judge (1970) reports a broad hornblende province located both north and south of San Francisco Bay in central California, and smaller epidote provinces located on either side of Point Conception in southern California. Clemens and Komar (1988) found a general northward dispersal of garnet in Oregon beaches, resembling similar patterns reported for offshore shelf deposits (Schiedegger et al., 1971). Heavymineral studies generally do not establish net littoral transport directions within individual littoral cells. The lack of specificity arises from 1) gradational changes in provenance mineralogy along the coast, and 2) mixing of provenance trace minerals on the continental shelf during late Pleistocene times. By comparison, dune field positions might reflect long-term directional wave and current forcing in discrete littoral cells.

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3. Methods The areal extents of the Holocene dune fields are taken from published maps (Cooper, 1958, 1967; Reckendorf, 1975; scale 1:40,000), NASA satellite images, and airphotos (scale 1:24,000). The map and airphoto data have been groundtruthed with auger and roadcut sections profiled at 650 sites (Peterson et al., 2006; Knott and Eley, 2006). Dune field midpoints are taken at the shoreline at equal shoreline distances between opposite ends of the mapped Holocene dune fields. Dune topsoil chronosequences, radiocarbon, and luminescence dating verify the Holocene ages of the dune fields presented here, as detailed elsewhere (Peterson et al., 2006; Knott and Eley, 2006). Cooper (1958, 1967) argued that the dune field positions might be controlled by sea cliff height, where low sea cliffs permit landward dune migration. To test this argument the average heights of mid-Holocene sea cliffs under the Holocene dune ramps are compiled from published sea-cliff cross-sections (elevation ±1.5 m MSL) (Peterson et al., 2006). In areas where dune cover does not reveal marine terrace height, the nearest uncovered marine terrace height is used (elevation ± 1.5 m MSL), as shown in US Geological Survey Seamless Maps (U.S. Geological Survey, 2007). Published boundaries of littoral cells, as mapped by Stembridge (1975), Swartz et al. (1985), Peterson et al. (1994), and Masters (2006), are used for the study region. Specifically, a littoral cell is defined here as a beach-fronted shoreline that is interrupted by rocky shorelines, often

Fig. 2. Diagram of a hypothetical littoral cell shows measured headland projection, shoreline angles, shoreline inflection points, shoreline segments, and dune field midpoint. Headland projections are measured perpendicular to adjacent shoreline trends, corresponding to either north or south bounding conditions. Shoreline angles, as measured relative to true north (TN) at 0.5 km intervals (Peterson et al., 1994) are averaged (mean) for north, central and south segments of the beach-fronted shorelines. The north, central, and south segments are defined by major inflections of shoreline angle. Averaged shoreline angles for the littoral shoreline segments are reported in negative degrees west of north and positive degrees east of north.

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headlands, of at least 1.0 km in alongshore distance parallel to the beach trend (Fig. 2). The concept of a closed littoral cell with distinct river sources and submarine canyon sinks (Inman and Frautschy, 1965) does not apply to most of the study area beaches (Peterson et al., 1991). Sand mixing within the littoral cells is assumed to occur within inter-annual to inter-decadal timescales. The magnitudes and timescales of sand transport that occur between some littoral cells via headland bypassing is a topic of current research in the study area. Headland projection distance and shoreline angles (Fig. 2) are measured from 1:24,000 scale topographic maps (U.S. Geological Survey, 2007) following Peterson et al. (1994). Shoreline angles are measured by bearing in degrees from true north (TN), negative values represent west of true north and positive values east of true north (Fig. 2). Shoreline angles are averaged for northern, central, and southern segments of littoral cells.

The averaging of shoreline angles (taken at about 0.5 km intervals) reduces bias from localized cusps, bars, rip-cells, sea-stacks, and creek mouths. The extents of the north, central, and south shoreline segments are based on major changes (inflections) in average orientation of the beach-fronted shorelines. The shoreline angles of the narrow beaches are not expected to have changed substantially relative to the regional scales of paleoclimate modeling in late Holocene times (5–0 ka). Resistant headlands and/or wave-cut platforms that are cut into bedrock help to armor the shoreline geometries in this uplifted coastline. Bathymetric contours of the innershelf are generally shore-parallel within the littoral cells. For this regional study, the shoreline angle is used as the proxy for local variation in wave incidence angle. Wave refraction modeling might refine local variation in wave angle, but should not reverse the general north or

Fig. 3. Maps of dune fields (shaded gray with four letter code), littoral cells (numbered 1–20), named headlands (horizontal bars), shoreline segments (lines), and asymmetry of dune field position (arrows) from (a) Point Grenville, Washington to Cape Blanco, Oregon, (b) Cape Blanco to Point Arena, and (c) Point Arena to Point Arguello, California. Letters designate sequential coastline sections from A (north) to U (south) as shown in the regional map (Fig. 1). Dune field and cell properties are presented in Tables 1 and 2, respectively. Headland position coordinates in figure parts a, b, and c, are given in UTM nothings (meters). Arrows show directional asymmetry of dune field positions within headland-bounded littoral cells. Dune fields (NEWP, FLOR, COOS, BAND) that are located in the center of corresponding littoral cells in central Oregon (Fig. 3a) do not show asymmetry of dune field distribution (double terminated lines). All other dune fields show directional asymmetry of position relative to cell midpoints (arrows).

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Fig. 3 (continued).

south direction of dominant wave approach in the study area littoral cells. 4. Compilation results 4.1. Coastal dune distribution The US West Coast dune fields are spatially discrete, with the important exceptions of the large dune sheets in the Columbia River littoral cell (CRLC) and in the south-central Oregon coast (Fig. 3). The dune fields range from a few kilometers to several tens of kilometers in along-coast length. The smaller dune fields are generally separated by long stretches of cliff-backed coastline with narrow beaches. The small dune fields range from 2 to 30 km2 in surface area (Table 1). The combined surface areas of the prograded barriers and beach plains in the Columbia River littoral cell total 208 km2 (Fig. 3a). The large size of these lowland dune fields reflects a very large supply of littoral sand from the Columbia River in Holocene times (Herb, 2000).

Large upland dune fields are developed in south-central Oregon, including FLOR (~73 km2) and COOS (~ 57 km2). These dune fields were fed by the shoreward wave transport of shelf sand from an offshore depocenter during the Holocene transgression (Peterson et al., 2007). A large upland dune sheet SANF (~75 km− 2) covers much of San Francisco City area in northern California (Fig. 3b). Submergence of San Francisco Bay during the Holocene transgression trapped coarse fluvial sediments in the Sacramento–American–San Joaquin bay-head delta (Atwater and Belknap, 1980). The most likely source of sand supply to the SANF dune sheet was shoreward sand transport from an offshore depocenter. Such a depocenter could have developed in the broad shelf located offshore of San Francisco Bay prior to the Holocene transgression. 4.2. Holocene dune development The earliest ages of upland dune development are provided by 1) available radiocarbon and luminescence dates from basal units of the upland dune fields (Orme and Tchakerian, 1986; Orme, 1992; Peterson

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Fig. 3 (continued).

et al., 2006; Knott and Eley, 2006), and 2) a reported basal date from a barrage lake formed behind Holocene dunes in the BAND dune field (Kelsey et al., 2005) (Table 1). Dates for 14 basal deposits from the upland dune fields range from approximately 8 to 4 ka. Holocene ages are established for undated dune fields, including CROO, CRES, ARCA, TENM, PTARE, BODE and PTRE, based on 1) weakly developed soil profiles in the dune fields, and 2) a lack of interbedded loess layers in the dune sand strata (Peterson et al., 2006). The earliest dated beach deposit in Oregon is from cell # 5. A clamshell from the top of a basal gravelly-sand lag (FLOR25a), located ~ 150 m landward of the modern beach (UTM-N 4867580), yielded a marine reservoir corrected date of 8381–8412 cal 14C yr BP ± 1σ (Beta-185769) (Peterson et al., 2006). Pismo beach clams (Tivela stultorum) from Native American middens in southern California, suggest the onset of beach development just north of the GUAD dune field at 9 ka (Masters 2006). However, most basal Tivela dates in the

southern California middens indicate that the earliest beaches formed at 7–5 ka. The onset of net beach accretion in the Columbia River littoral cell (Fig. 3a) ranges from 2 to 4 ka (Rankin, 1983; Woxell, 1998). The largest, upland dune fields in the study area generally began to develop (8–7 ka) soon after the onset of beach sand supply at about 9–8 ka (Table 1). Some smaller dune fields did not begin development until well into the late Holocene high stand at 5–4 ka. The several thousand year time lag for these smaller dune fields to develop might indicate that a period of longhore sand transport and beach sand redistribution were needed to initiate upland dune accumulation. The apparent delay in the onset of the lowland dunes in the Columbia River littoral cell reflects the initial filling of shelf- and bay accommodation space prior to net beach accretion there (Herb, 2000). Radiocarbon ages from dune field tops range from 0.5 to 0.1 ka (Table 1). Nearly all of the study area dune fields showed some historic dune mobility prior to artificial stabilization during the

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Table 1 Holocene dune fields in the U.S. West Coast of North America. Dune field name

Midpoint position (UTM)

N–S length (km)

Backedge (m)

Ave. width (km)

Surf. area (km2)

Basal/top dune dates (ka)

Basal dune elev. (m)

CRLC

5160900

104

5

2.0

208

5

MANZ NETA4 PACI NEWP FLOR COOS BAND GOLD8b CROO CRES ARCA TENM PTAR BODE PTRE SANF ANON MONT PSUR MORR GUAD SANT

5062840 5033240 5017275 4932760 4863100 4820620 4773055 4702210 4679105 4632905 4527090 4375340 4313960 4242780 4213980 4176430 4108330 4058920 4019850 3908370 3877320 3854680

4.0 – 12.5 5.5 38.5 27 23.5 – 3.6 5.4 8.8 7.1 5.6 3.9 15.8 12.5 2.8 22 3 4.3 18.3 11.5

26 – 64 25 27 39 14 – 43 10 8 24 15 20 25 86 15 16 20 32 43 29

1.1 – 1.5 0.3 1.9 2.1 1.3 – 0.9 1.9 0.6 0.9 0.7 0.8 0.5 6.0 0.6 0.3 0.3 0.6 1.5 1.4

4.4 – 19 2.0 73 57 30 – 3.2 21 5.3 6.4 3.9 3.1 7.9 75 1.7 6.6 0.9 2.6 27 15

~ 4.2/~0.1 ~ 2.51/~ 0.1 –/~ 0.3 ~ 6.5a/– ~ 7.2/~ 0.2 ~ 6.5a/~ 0.1 ~ 7.7/~0.3 ~ 5.6/– ~ 7.02/~ 0.4 ~ 5.9a/– –/~ 0.5 – – – – – – ~ 7.4/– ~ 6/– ~ 8/– – ~ 63/– ~ 4.34/–

10 – 10 15 – – 15 – 10 10 5 10 20 15 15 10 10 10 15 20 10 15

Dune field names from Peterson et al. (2006) include lowland dune fields in the Columbia River Littoral Cell (CRLC) and upland dune fields including Manzanita (MANZ), Pacific City (PACI), Newport (NEWP), Florence (FLOR), Coos Bay (COOS), Bandon (BAND), Crook Point (CROO), Crescent City (CRES), Arcata (ARCA), Ten Mile (TENM), Point Arena (PTAR), Bodega (BODE), Point Arena (PTRE), San Francisco (SANF), Ano Nueva (ANON), Monterey (MONT), Point Sur (PSUR), Moro Bay (MORO), Guadalupe (GUAD) and Point Sal (PSAL), and San Antonio (SANT). Dune field midpoint (see Fig. 2) located by UTM-Northing coordinate. Backedge elevation: average of 10–50 dune ridge or backedge elevations (m) relative to mean sea level (MSL). Average width: average of 10–50 dune widths taken perpendicular to shoreline. Dune basal dates are from 1: Woxell (1998), 2: Kelsey et al. (2005), 3: Orme and Tchakerian (1986), 4: Knott and Eley (2006). All remaining dates are from Peterson et al. (2006, 2007). Dune basal elevation (m MSL) is an average taken from at least 5 exposed sea cliff sections (Peterson et al., 2006) or from uncovered terrace tops taken immediately adjacent to dune fields from U.S. Geological Survey Seamless Map Data (U.S. Geological Survey, 2007). a Basal Dates: Thousand years (ka), retreat scarp (RS), radiocarbon (RC), thermoluminesence (TL), barrage lake bottm (BL), and basal dune date from adjacent, small dune field.

middle 1900s (Cooper, 1958; 1967). Sand supply to the dune fields continued throughout the latest Holocene. However, wave truncation of upland dune ramps and/or recent erosion of wave-cut platforms in Oregon (Hart and Peterson, 2007) suggest that sand supplies to many beaches, and associated dune fields, are now in decline. 4.3. Littoral cells With the exceptions of barriers and beach plains associated with the Columbia River (cell #1) the littoral cells in the study area (5–86 km in north–south length) are separated by rocky shorelines or headlands (Fig. 3). Twenty littoral cells are identified with corresponding dune fields in the study area (Table 2). The littoral cell boundaries are identified as rocky headlands (less than 10 m beach width) of at least 1 km in length (Peterson et al., 1991; 1994). Littoral cells that lack corresponding dune fields are not addressed in this paper. Those littoral cells without dune fields account for about two thirds of the shoreline length in the study area (Fig. 3).

4.4. Sea cliff height In late Holocene times, the upland dune fields ramped over sea cliffs that had previously been cut into marine terraces during the mid-Holocene transgression. Elevations of the sea-cliff terraces that were overtopped by the dunes reach as much as 10–20 m above mean sea-level (basal dune elevations from Table 2). As indicated by Cooper (1958, 1967) upland dune development appears to be precluded by the highest sea cliffs. Of importance to this study is the potential control of sea-cliff height on dune field position within the littoral cells. Generally, low rates of tectonic uplift (0–0.2 mm yr− 1) (West

and McCrumb, 1988; Muhs et al., 1990; Lettis and Hall, 1994) have yielded modest terrace elevations of 5–20 m MSL in the larger cells. We test the potential control of sea-cliff height on dune field position by comparing the relative positions of dune fields and lowest terrace surface (Tables 1 and 2). The dune field midpoints do not generally correspond to the positions of the lowest terrace surfaces (Fig. 4). Separation distances between the lowest sea cliffs and dune field midpoints range from 1 to 25 km, and average 9 km for the 20 littoral cells. Basal dune elevations and lowest Pleistocene terrace surface elevations are compared to test potential control of sea-cliff height on dune field position. Most of the dune fields that overtopped sea cliffs are 30– 50% higher than the corresponding lowest terrace elevations in corresponding cells (Fig. 5). The averaged backedge elevations of the upland dune fields reach 25 to 80 m MSL, which is 2–5 times higher than the corresponding sea cliffs that were overtopped by the dune fields. Given an excess beach sand supply, there has been sufficient onshore wind power to mobilize dune fields well above the sea cliffs in the analyzed littoral cells. Although sea-cliff height is not the primary control of upland dune field development in the littoral cells investigated here it is likely to be important in shorelines with high sea cliffs and low rates of beach sand supply. 4.5. Dune field position within littoral cells Dune field midpoints are evaluated with respect to relative position (north, central or south) within corresponding littoral cells (Fig. 3). The cell shoreline segments as divided by inflections in shoreline angle (Fig. 2; Table 2). The relative positions of dune fields in their hosting cells change from north to south in the study area (Fig. 6). Dune fields in northernmost Oregon occur at the north ends of their littoral cells. Dune fields in southernmost Oregon and California, generally occur at the

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Table 2 Littoral cells containing dune fields in the U.S. West Coast of North America. Cell #/dune name 1

N. head UTM-N/projection S. head UTM-S/projection North inflection UTM(km) (km) N/angle (°)

5239500 (0.6) 2 5065800 MANZ (1.3) 3 5020800 PACI (2.7) 4 4947500 NEWP (1.5) 5 4885900 FLOR COOS (0.4) 6 4795900 BAND (2.5) 7 4686500 CROO (2.5) 8 4655300 CRES (1.6) 9 4545000 ARCA (3.5) 10 4383600 TENM (0.6) 11 4321200 PTAR (0.6) 12 4257700 BODE (1.0) 13 4241500 PTRE (2.7) 14 4187500 SANF (0.9) 15 4115356 ANON (0.9) 16 4092300 MONT (1.4) 17 4024000 PSUR (0.9) 3926700 18 MORR (3.2) 19 3893100 GUAD (2.2) 20 3864600 SANT (2.2) CRLC

5090000 (2.5) 5037700 (0.7) 4991700 (1.3) 4906700 (0.4) 4799700 (0.8) 4748000 (0.5) 4678500 (0.3) 4626000 (0.6) 4485000 (0.3) 4371400 (1.4) 4312200 (0.4) 4241500 (0.4) 4205400 (0.8) 4161000 (2.0) 4107200 (0.5)a 4053300 (2.5) 4018600 (0.3) 3905900 (0.9) 3864600 (0.3) 3848600 (0.5)

5222500 (− 10) 5063900 (− 24) 5017300 (− 15) 4945600 (− 6) 4841300 (0) 4788100 (− 6) 4684300 (− 20) 4650600 (− 40) 4541600 (− 25) 4380300 (− 19) 4320800 (− 4) 4254400 (− 40) 4225100 (− 20) 4181400 (− 33) 4162800 (− 44) 4074300 (− 25) 4022500 (− 22) 3921500 (− 30) 3892200 (− 40) 3860200 (− 43)

Central shoreline UTMN/angle (°)

South inflection UTMN/angle (°)

Lowest terrace UTM-N/ elev. (m)

5160900 (− 4) 5052100 (+ 4) 5008900 (+ 2) 4930500 (+ 2) 4824500 (+ 6) 4772300 (+ 9) 4682600 (− 1) 4644300 (0) 4533800 (+ 10) 4379300 (10) 4317800 (+ 16) 4248800 (− 25) 4223300 (+ 02) 4172100 (− 8) 4136100 (+ 2) 4068100 (+ 4) 4021900 (− 1) 3916500 (+ 5) 3886600 (+ 7) 3857200 (+ 6)

5099300 (+7) 5040300 (+10) 5000600 (+9) 4913548 (+6) 4807800 (+24) 4756600 (+27) 4680900 (10) 4638000 (+14) 4526000 (+27) 4378400 (+15) 4314900 (+28) 4243200 (19) 4221400 (+20) 4162800 (+43) 4109500 (+12) 4061800 (+29) 4021400 (+19) 3911600 (+19) 3881100 (+17) 3854300 (+20)

– 5088200 (5) 5011500 (5) 4914930 (6) 4871400 (5) 4826400 (7) 4683100 (8) 4647100 (7) 4541600 (5) 4375700 (7) 4316600 (8) 4244300 (6) 4221900 (15) 4180106 (8) 4110200 (5) 4074800 (5) 4020100 (5) 3922200 (8) 3887400 (5) 3855300 (8)

Cell number and corresponding bounding headland names including (1) Point Grenville–Tillamook Head; (2) Neahkanie Mt–Cape Mears; (3) Cape Lookout–Cascade Head; (4) Yaquina Head–Yachats; (5) Sea Lion Point–Point Gregory; (6) Cape Arago–Blacklock Point; (7) Humbug Mountain–Crook Point; (8) Point Checto–Point St George; (9) Trinadad Head–False Cape; (10) Point Bruel –Point Laguna; (11) Point Irish–Point Arena; (12) Point Duncan–Mussel Point; (13) Mussel Point–Point Reyes; (14) Point Tennessee–Point San Pedro; (15) Pigeon Point–Ano Nuevo Island; (16) Capitola Bluffs–Point Cabrillo; (17) Point Hurricane–Point Sur; (18) Point Estero–Spooners; (19) Point San Luis–Point Sal; (20) Point Sal–Point Purisima. Dune field names from Table 1. Northern and southern bounding headlands (N. Head and S. Head) with UTM-Northing positions and seaward projection distance (km), see Fig. 2 for details. North and south inflection points (see Fig. 2) include UTM-Northing positions, and average shoreline angle, as given in degrees west (−) or east (+) from true north (TN). Lowest terrace height is taken either from measured exposed sea cliffs, (Peterson et al., 1991; Peterson et al., 2006) or Seamless Map Data with 3 m supplement contours, yielding uncertainty of ± 1.5 m (U.S. Geological Survey, 2007). a Año Nuevo headland projection is estimated where a breach of the paleoheadland produced Año Nuevo Island.

south ends of their cells. Dune fields in central Oregon are either centrally located or extend along the full length of their littoral cells. Three dune fields, CRLC (cell #1), ARCA (cell # 9), and SANF (cell # 14), do not conform to the general pattern of dune field position in the study area (Fig. 6). The anomalous positions of the CRLC and ARCA dune fields are related to exceptional shoreline angles. The anomalous position of the SANF dune field in littoral cell # 14 is not due to local shoreline angle because shoreline angle in that cell is essentially north–south (Fig. 3c). Some factor other than regional wave approach angle is therefore controlling net littoral drift in the shoreline south of the San Francisco Bay mouth. 4.6. Shoreline angle Fig. 4. Plot of separation (north–south distance in km) between shoreline locations of dune field midpoint and lowest terrace surfaces for 19 littoral cells (numbered 2–20 from north to south) in the study area. Two dune fields (FLOR and COOS) are shown for littoral cell 5 (see Fig. 3 for maps of cell and dune field locations).

Within the relatively straight coastlines of the study area there are significant local variations in shoreline angle (Fig. 3). With only two exceptions (cells # 11 and # 12) the central shoreline segments are within

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Fig. 5. Plot of upland dune field heights and corresponding terrace elevations in littoral cells, numbered 1–20 from north to south. Dune terrace elevations (m MSL) are from basal dune elevations (Table 1) as taken from average heights of cell terrace surfaces overtopped by corresponding dune fields. Dune backedge elevations (m MSL) are from averaged elevations (Table 1) of dune backedges or highest ridges in the dune field. Lowest terrace elevations (m MSL) in the littoral cells are taken from Table 2.

10° of a north–south orientation. By comparison, the northern shoreline segments average 24° west of north, and the southern shoreline segments average 19° east of north (Table 2). The large shoreline angles near the headlands locally decrease the angle of oblique wave attack, as shown in Fig. 7A,B. For example, oblique attack angle diminishes from the south at a northern headland in cell # 2 located in northern Oregon, and from the north at a southern headland in cell # 8 located in northernmost California. Longshore currents diminish with decreasing wave incidence angle (Komar, 1979) thereby reducing longshore transport. The localized, and probably episodic,

accumulation of beach sand in the longshore convergence zones feeds the coastal dune fields. Longshore transport trends can locally reverse in direction due to opposing shoreline angles. For example, littoral cells # 1 and # 9 demonstrate unexpected, centrally distributed dune fields that result from anomalous shoreline angles (Fig. 7C,D). An opposing shoreline angle in the southern end of cell #1 displaces some CRLC sand accumulation to the south of the Columbia River in the Clatsop Plains. High shoreline angles at the northern end of the cell preclude substantial sand accumulation in the northern beaches. In contrast, the northeast-trending shoreline in cell # 9 displaces the ARCA dune field to the north of the Eel River mouth. This position is reversed from the regional pattern expected for the California littoral cells (Fig. 6). The shoreline angles within the sea-cliff-backed littoral cells are related to differential erosion of the resistant headlands and the less resistant embayments over successive marine high stands (Inman and Frautschy, 1965; Muhs et al., 1990; Peterson et al., 1991). Headland projections, normal to the adjacent shorelines, range from 0.3 to 3.5 km in distance. Fourteen littoral cells show asymmetry of dune field positions (Fig. 6). In those 14 cells the dune field midpoints occur between the nearest shoreline inflection point and the associated bounding headland. The relative importance of these two features in controlling littoral sand entrapment is addressed by comparing their separation distances from the dune field midpoints. Generally, the dune field midpoints are closer to the shoreline inflection points, average 2.3 km separation distance, than to the bounding headlands, average 4.2 km separation distance (Fig. 8). Wave refraction at the headlands can also contribute to the apparent longshore flow convergence represented by the dune fields.

4.7. Headlands

Fig. 6. Plot of dune field position in north, central, south shoreline segments of corresponding littoral cells in the study area. Northing coordinates (UTM-N) and state boundaries WA, OR, CA are shown for reference to study area latitudes (Fig. 3). Littoral cells are numbered (#1–20) from north to south. Dune field names are shown for each littoral cell. Dune field midpoint positions and cell segment data are from Tables 1 and 2.

The role of the headlands in limiting littoral transport is unclear. There is no apparent correlation between headland projection distance and dune field size in the 14 littoral cells that show dune field position asymmetry. Most of the littoral cells in the study area do not include upland dune fields (Fig. 3). Either these cells lacked sufficient sand supply during the late Holocene or their littoral sand was lost to other sinks, such as adjacent cells. Sand bypassing around small headlands is

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tested between littoral cells where regional net longshore transport is established by asymmetric dune field positions in nearby cells. In northern Oregon, two adjacent littoral cells demonstrate widely different dune field development (Fig. 9). The southern cell (Government Point–Lands End) shows narrow, retreating beaches and no upland dune development. The northern cell (# 3) contains a large headland, Cape Lookout (2.7 km projection) and a substantial upland dune field PACI (19 km2) at its northern terminus. These paired littoral cells are similar in size, shoreline angles, and associated small tidal basins (Peterson et al. 1991). The very small headlands at Lands End and Cascade Head (0.4 km projection) permitted the bypassing of sand from the southern cell to the northern cell. The large Cape Lookout headland apparently serves as a northern boundary to some sand transport in cell # 3.

The paired littoral cells in southern Oregon show a similar but opposite directional trend in headland bypassing (Fig. 9). The Rogue River (Fig. 1) yields a significant bedload to the associated littoral cell (Humbug Mt– Cape Sebastian; Peterson et al., 1991). A small dune ramp delivered sand to Otter Point (~5.9 ka; Table 1) but no upland dune fields occur in that cell (Fig. 3b). In contrast, the southern cell (# 7) includes an upland dune field (CROO) at its southern terminus. The CROO dune field accumulated surplus littoral sand through latest Holocene times (top dune date 0.5 ka; Table 1). The small headland separating the paired cells, Cape Sebastian (0.3 km projection), permitted bypassing of Rogue River sand to the south. The headland at the southern end of the southern cell, Crook Point, is also small (0.3 km projection). The beach sand entrapment feeding the CROO dune field results from shoreline angle change rather than headland

}

Fig. 7. Maps of selected littoral cells showing bounding headlands, beach-fronted shoreline, shoreline angles, and dune field positions. In cell # 2 (A) the dominant wave direction is from the southwest and the predominant littoral transport is to the north. In cell # 8 (B) the dominant wave direction is from the northwest and predominant littoral transport is to the south. Northwest-trending shorelines in cell #1 (C) displace Clatsop and North Beaches sand accumulation to the south. Anomalous northeast-trending shorelines in cell #9 (Part D) displace the upland dune field ARCA to the north.

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Fig. 7 (continued).

projection. The relatively small size of the CROO dune field (3.2 km2) and its very close proximity to the small headland (0.3 km separation distance; Fig. 9) suggest that some littoral sand is bypassing Crook Point to the south. More work is needed to establish the trapping efficiencies of headlands and associated shoreline angles in this high wave energy regime (see Section 2).

south-central Oregon started to develop in latest mid-Holocene times (Peterson et al., 2007). Dated basal strata from seven of the smaller, upland dune fields indicate that sand accumulation began at 6–4 ka (Table 1). Near-surface deposits in the dune fields date to within the last few centuries or modern times. Sand accumulation in the coastal dunes continued throughout the late Holocene, though possibly at diminished rates in more recent times (Hart and Peterson, 2007).

5. Discussion 5.1. Paleoclimate models Net littoral transport directions are difficult to establish in littoral cells with narrow or eroding beaches that experience seasonal reversals in wave and current directions. In such cells only the upland dunes are preserved over long time periods to serve as proxies for net littoral transport. Strong asymmetry of dune field distribution is observed in 11 littoral cells of the US West Coast (Fig. 6). The largest upland dune fields in

We use simulations from a regional climate model for time intervals at 11, 6 and 0 ka to compare regional wind fields and associated wave forcing with Holocene dune field proxies of net littoral transport. The modeling approach, which nests a regional climate model (RegCM2) with an atmospheric general circulation model (AGCM, GENESIS, V2.3),

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Fig. 8. Plot of separation distance between the dune field midpoints and the nearest shoreline inflection point (solid square) and headland (solid hexagon) in 14 littoral cells showing asymmetry of dune field position. Dune fields are generally closer to shoreline inflection points than to headlands.

has been used in several paleoclimate studies (Hostetler et al, 1994; Hostetler and Bartlein, 1999; Hostetler et al, 2000). For this study, timevarying boundary conditions for the regional model were produced from 50-year simulations with GENESIS for each time slice. The RegCM2 simulations are 10years long, and the analyses are done on the last 8 years of simulation intervals to allow for model spin-up. This approach was followed, using a different domain and version of our paleomodel, for late Pleistocene periods (30–21 ka) of directional wave forcing and low-stand dune development (Peterson et al., 2007). To accommodate potential forcing from wide-ranging wave generation and local wind stress, the paleoclimate conditions within 0 to 2000 km of the coast were targeted. The most extreme directional wave forcing from oblique wave attack angle was expected to come from storm centers that are closest to the coast. We assumed that deepwater wave forcing and innershelf currents are generally related to seasonal wind directions, with winter and summer months yielding the greatest differences. Seasonal patterns of sea-level pressure are associated with seasonal patterns of wind strength and direction (Fig. 10a). During winter months the low pressure (995 mbar) that is associated with the Aleutian Low in the Gulf of Alaska creates anticlockwise (cyclonic) airflow that results in southwest winds along much of the Oregon and Washington coasts. Those southwest winter winds generate large surface waves that travel toward the northeast. Those winds also drive geostrophicallytrapped currents that travel northward on the innershelf (Sternberg, 1986). By comparison, winter wind flow is generally perpendicular to the northern California coast, and it approaches the southern California coast from the northwest. The northwest winter winds drive waves towards the southeast in the central and southern California coasts (Fig. 11). Summer wind directions that are associated with the land–sea temperature contrasts and the subtropical-high are uniformly directed from the north–northwest. The summer winds drive small fair-weather waves towards the southeast in Washington, Oregon, and California.

Spring and autumn wind forcings are intermediate between the winter and summer extremes. Modeling results show the low-pressure center in the northeast Pacific to extend from the Gulf of Alaska to southern Oregon during the winter months of early Holocene times at 11 ka (Fig. 10b). The offshore trough of low pressure (simulated average value 1005 mbar) was not strong compared to present conditions (average 995 mbar). However, it combined with anticyclonic winds associated with surface high pressure from the upper continental interior to drive coastal winter winds from the south in Washington, Oregon, and northernmost California. Northward littoral transport during the period of winter storms should have dominated the paleo-coastline from Washington to central California at 11 ka (Fig. 11). Summer winds would have been from the north along the entire study area coastline at this time. The coastline at 11 ka was located ~5 to 20 km offshore, in the present innershelf, when the corresponding sea level was 50–60 m below present sea level (Peterson et al., 2007). At 6 ka the winter offshore low-pressure trough deepened to a simulated average value of ~995mbar, creating strong winds and waves, but its position was still centered between the Gulf of Alaska and Washington State (Fig. 10c). Winter wind flow and waves approached the Washington and Oregon coasts from the south and southwest, respectively. Continental high-pressure winds diminished substantially, permitting winter winds and waves to approach the northern California coast from the west. Summer winds at 6 ka are modeled to strike the coast from the north along the entire study area. The strong onshore wind and wave forcing during winter months in mid-Holocene time would have facilitated onshore transport of existing shelf sand deposits. The innershelf was largely submerged by 6 ka when sea level was only ~10 m below present sea level (Peterson et al., 2007). Winter waves from the southwest continued to drive northward littoral transport in Washington and Oregon at 6 ka, but northern California might not have experienced a predominant littoral drift direction in mid-Holocene times. During early- to mid-Holocene time, an intensification of the offshore winter trough of low pressure and diminishing high pressure in the continental interior altered littoral transport directions in the US West Coast. Strong northward wave forcing in Oregon and northern California during earliest Holocene times underwent a transition to eastward (onshore) forcing by mid-Holocene times. By late Holocene times paleoclimate conditions yielded divergent wind and wave directions in the 1500 km long study area. Late Holocene winter wind and wave forcing generally approached Washington from the southwest, Oregon from the west, and California from the northwest (Fig. 11). The distributions of late Holocene dune fields (5–0 ka) in littoral cell shoreline segments (Fig. 6) are in general agreement with modeled winter paleoclimate conditions interpolated between 6 and 0 ka. 5.2. Heavy-mineral tracers of littoral transport Specific heavy minerals in some beach deposits of the study area provide constraints on net littoral drift over broad areas. For example, an orthopyroxene mineral (hypersthene) characterizes Holocene sand supply from the Columbia River mouth in cell # 1 (Fig. 12A) (Baker, 2002). The outer shelf was supplied with a clinopyroxene mineral (augite) in late Pleistocene times. The dominance of hypersthene relative to augite in the innermost shelf of northern Washington (Venkatarathnam and McManus, 1973) demonstrates a northward, net littoral transport that extends at least 200 km north of the Columbia River mouth. Although some shorelines in cell #1 indicate localized southward transport (Fig. 7C), the broader innershelf transport is to the north in Washington. Northward bypassing of Columbia River sand around Point Grenville and Hoh Head is probably facilitated by combined nearshore and shelf geostrophic flow (Fig. 11). In northern California, the Klamath River sand is characterized by bluegreen hornblende, whereas the Eel River is characterized by glaucophane (Fig. 12B). Both the Klamath and Eel Rivers deliver significant bedload to

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Fig. 9. Two paired littoral cells from northern and southern Oregon showing sand bypassing around small headlands at Lands End-Cascade Head (Part A) and Cape Sebastian and possibly Crook Point (Part B).

the littoral system. The distinctive mineral glaucophane is locally traced north of the Eel River mouth in the high-angle shoreline of cell # 9 (Bodin, 1982) confirming northward beach transport to the ARCA dune field (Fig. 7D). However, blue-green hornblende occurs in beach sand located at least 100 km south of the Klamath River, denoting southward net littoral drift around Trinidad Head in late Holocene times. The southward transport of blue-green hornblende during the late Holocene contrasts with a general northward dispersal of glaucophane during late Pleistocene times (Schiedegger et al., 1971). The reversal of net littoral transport from late Pleistocene to late Holocene time is consistent with modeled changes in wind and wave forcing during the Holocene (Fig. 11). Green–brown hornblende dominates the heavy-mineral fraction of beach sand in San Francisco as well as in beach sand located as far

north as Point Reyes and as far south as Halfmoon Bay and Año Nuevo (Figs. 1, 3c, and 12C) (Judge, 1970). However, the SANF dune field (Bonilla, 1971) is restricted to a position immediately south of the Golden Gate, the entrance to San Francisco Bay. Local northward littoral transport might result from anomalous tidal–current effects near the Golden Gate. Opposing giant sand waves in the Golden Gate (U.S. Geological Survey, 2006) represent a counter-clockwise tidal gyre that could generate net northward transport in beaches located south of the Golden Gate. The regional pattern of green–brown hornblende dispersal in beaches both to the north and south of San Francisco Bay (Judge, 1970) is not consistent with southward transport shown by dune fields in cells # 12, 13 and 15 (Fig. 3c). Some northward transport must have

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occurred on the broad continental shelf offshore of the Golden Gate during the late Pleistocene or earliest Holocene, prior to net southward forcing predicted for late Holocene times (Fig. 11). In southern California, a decline in epidote abundance with distance south of Pismo Beach suggests limited sand transport south to Point Conception (Fig. 12D) (Judge, 1970). Small dune fields located north of Point Arguello and at Point Conception confirm some southward littoral transport to Point Conception. The ages of the SANT dunes are reported

to be late Holocene (Chambers Consultants and Planners, 1984). Initiation of the northern GUAD dune sheet at ~4.3 ka (Table 2) substantially post-dates beach sand accumulation to the north at Pismo Beach (7–9 ka; Masters, 2006). Sand transport in the most southern cell # 20 (Table 2) must have reversed between early and late Holocene times. The dune positions, onset ages of dune development, and epidote abundance all reflect southward net littoral drift at the southern end of the study area in late Holocene times (Fig. 11).

Fig. 10. a. Seasonal averages of sea-level pressure (SLP) in millibars (mb) and associated surface–wind directions (arrows) for modern conditions (0 ka). Wind strength (shown by arrow size) is proportional to pressure gradient (Patrick and Hostetler, 2004). Model results shown here are from updated runs completed in 2007. b. Seasonal averages of sea-level pressure (SLP in mb) and associated surface–wind directions (arrows) for earliest Holocene conditions (11 ka). Arrow size is relatively proportional to wind strength. c: Seasonal averages of sea-level pressure (SLP in mb) and associated surface–wind directions (arrows) for mid-Holocene conditions (6 ka). Arrow size is relatively proportional to wind strength.

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Fig. 10 (continued).

5.3. Regional trends of net littoral transport in the U.S. West Coast Net littoral transport within a specific shoreline segment is the product of variable wave climate, innershelf currents, innershelf bathymetry, and local shoreline angles. However, regional trends of net littoral transport should emerge over long distances of relatively straight coastline. The positions of late Holocene dune fields in littoral cells (Fig. 6) and reported patterns of heavy-mineral tracers in four littoral systems (Fig. 12) are compiled to establish regional patterns of predominant littoral transport direction in the study area (Fig. 13).

Net northward transport predominates from Government Point in northern Oregon to northern Washington, with a short reversal just south of the Columbia River mouth. There is no apparent net transport direction from Government Point to Cape Blanco in Oregon. Southward net transport is dominant in sandy beaches from Cape Blanco, Oregon to southern California, with two exceptions just north of Cape Mendocino and just south of San Francisco Bay. The late Holocene period of most dune development (6–0 ka) shows a strong correlation between expected wind–wave forcing (Fig. 11) and dune field indicators of predominant littoral transport in littoral cells (Fig. 13).

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Fig. 10 (continued).

The use of dune field distributions in littoral cells to establish longterm net littoral drift should be more widely used. Geomorphic evidence of net littoral drift is necessary to independently verify other means of evaluating net transport, such as wave-buoy data, wind– wave hindcasts, sediment provenance indicators, or tracer experiments. In accreting coasts the geomorphic indicators of net littoral drift, including downdrift shoreline accretion, bay spits, and migrating inlets, are well preserved. But such features are not preserved over long-term periods in marginally stable or eroding coasts. The high preservation potential of upland dunes provides a means of recording the relative residence time of sand in shorelines over a wide range of timescales. Given similar conditions of wind power, topographic

relief, and vegetation within littoral segments the variability of dune accretion alongshore provides long-term records of net sand transport to corresponding shorelines. These results should have broad application in other coastlines around the world that are characterized by stable or retreating shorelines, sandy beaches, and onshore winds. 6. Conclusions The longshore distribution of upland dune fields reflects residence time of beach sand in corresponding littoral cells of the US West Coast. Under conditions of variable wave climate and episodically eroding beaches the upland dunes represent long-term records of littoral transport

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Fig. 11. Summary diagram of modeled wind–wave forcing under winter and summer conditions along the US West Coast at 11 ka, 6 ka, and 0 ka. See Fig. 10 for modeled sea-level pressure and associated wind vectors at 11, 6 and 0 ka.

directions. Asymmetry of dune field development in the littoral cells demonstrates downdrift accumulation of beach sand against changes in shoreline angle and/or cell bounding headlands. Contrasting dune development in adjacent littoral cells that are supplied by a common sand source indicates sand bypassing around minor headlands and/or rocky shorelines. The timescales of net littoral transport are recorded in the dunes by dated accumulations of dune strata. The Holocene records of net littoral transport are consistent with modeled paleo-wind flows that drive nearshore wave approach and shelf geostrophic flow. Changing paleoclimate conditions resulted in a reversal of net littoral drift in the central study area from early Holocene to late Holocene times.

Acknowledgements This paper benefited from early reviews by John Dingler and Bob Morton. This research was funded by the NOAA Office of Sea Grant and Extramural Programs, US Department of Commerce, under grant number NA76RG0476, project number R/SD-04, and by appropriations made by the Oregon State Legislature. Unocal Corporation provided funding for the dating of dune deposits in the Guadalupe dune sheet of California. The US Geological Survey provided funding for the dating of barrier dune ridges in the Columbia River littoral cell of Washington.

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Fig. 12. ab: Heavy-mineral indicators confirm regional, northward net littoral transport in Washington (A), and regional, southward net littoral transport in northern California (B) in spite of local transport reversals in high-angle shorelines. cd: Heavy-mineral Indicators establish broad littoral provinces in the San Francisco Bay area and between Pismo Beach and Point Conception (D). Dune field distributions within shoreline segments of the large littoral systems denote 1) local transport reversal (northward) in San Francisco Beaches (C) and 2) southward transport towards Point Conception (D).

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Fig. 13. Generalized line drawing of study area with dominant littoral drift directions (arrows) based on dune field and tracer-mineral indicators of net littoral transport in representative littoral cells. Map coordinates are in Long and Lat (degrees) and UTM-northing (N) in 100 km intervals.

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