Geological Society of America Bulletincascadiageo.org/documentation/literature/cascadia... · rection and a time-varying gain correction for spherical spreading of the signal (cf

Geological Society of America Bulletin

doi: 10.1130/0016-7606(1998)110<0572:SAAAFO>2.3.CO;2 1998;110;572-587Geological Society of America Bulletin

Robert S. Yeats, LaVerne D. Kulm, Chris Goldfinger and Lisa C. McNeill Stonewall anticline: An active fold on the Oregon continental shelf

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ABSTRACT

Stonewall Bank is the site of a growingwest-verging anticline striking north-north-west on the continental shelf at 44.5° N,southwest of Newport, Oregon. To the eastare Pliocene-Pleistocene strata of the New-port syncline, onlapping eastward againstgently west-dipping late Miocene and olderrocks of the Oregon Coast Range. Folding ofStonewall anticline results in fine-grainedmiddle Miocene strata being exposed at thesea floor along the anticlinal crest; rocks asold as Eocene are encountered in the Unocal P-0093-1 Grebe well, drilled at the anticlinalcrest. Rates of folding are based on deforma-tion of an unconformity between Plioceneand Miocene strata (PM unconformity) andof a stream channel that crossed StonewallBank during the last glacial maximum. Thebed length of the unconformity is shortenedacross the Stonewall anticline and adjacentfolds by about 400 m between structures westof Stonewall Bank and the Newport syncline.The PM unconformity has a vertical separa-tion of about 1000 m between the anticlineand the first syncline to the west. The hori-zontal shortening and vertical separation im-ply that Stonewall anticline is underlain by ablind reverse fault. Retrodeforming the PMunconformity shows that this fault dips65°–70° E. A vertical separation of 1000 m ona fault with this dip yields a slip of 1070–1080m along the fault. If folding of the PM un-conformity is assumed to have begun 2–3Ma, this would give a long-term slip rate of0.4–0.6 mm/yr. If most folding began afterdeposition of the entire Pliocene-Pleistocenesequence, the slip rate would be 1.0–1.1

mm/yr. Stonewall anticline has arched thelate Pleistocene lowstand wave abrasion plat-form since sea level underwent a rapid risefrom 14.5 to 8 ka. This arch is crossed by anantecedent stream channel that is 275–550 mwide and is marked by side drainages andcut banks up to 12 m high. Warping of theplatform on the eastern limb of the anticlinehas back-tilted the stream channel eastwardtoward its present onshore continuation, theYaquina River. The platform slopes down-ward 10–13 m westward from the crest ofStonewall anticline. We estimate that theplatform stopped abrading when sea levelreached about –40 m at 11–12 ka. Assumingthat the west slope of the platform is con-trolled by the same blind fault that producedthe west dip of the PM unconformity, theHolocene slip rate on this fault would be0.9–1.3 mm/yr, comparable to the long-termslip rate. Rupture of the entire 25 km lengthof the blind fault with 1 m of slip beneathStonewall Bank could produce an earth-quake with Mw = 6.8 ± 0.25, which would re-sult in peak ground accelerations close to0.2g on the central Oregon coast.

INTRODUCTION

Deformed topographic features compriseone of the principal ways to measure tectonicstrain across active faults and folds on land(Keller and Pinter, 1996), but landforms thatrecord deformation on continental shelves arerelatively unstudied. For this reason, the earth-quake and tsunami hazards posed by offshorestructures are commonly unrecognized and are rarely quantified, although GeomatrixConsultants (1995) made an effort to do so forOregon. Continental shelves may be closer to aplate boundary than the better-studied onshore

regions, and slip rates may be higher on off-shore structures than onshore (Geomatrix Con-sultants, 1995). Such is the case for the Oregoncontinental shelf.

Sidescan sonar imagery, swath bathymetry,and submersible observations now allow themapping of deformed stream channels, waveabrasion platforms, and shorelines formed duringsea-level lowstands. Rates of deformation can bedetermined where these submerged landformscan be dated. Goldfinger et al. (1992; 1997a)used sidescan sonar to determine the Holoceneslip rate on several left-lateral strike-slip faults onthe continental slope and abyssal plain near theNorth America–Juan de Fuca plate boundary(Cascadia deformation front) off Oregon andWashington. Three of these faults are shown onFigure 1. Slip rates were measured from the off-set of dated submarine channels and of isopachsof Pleistocene abyssal-plain turbidites identifiedin multichannel seismic-reflection profiles acrossthe Astoria submarine fan, age-calibrated by adeep-sea core hole (Goldfinger et al., 1997a).

Here we compare rates of deformation of awarped Holocene stream channel with longer-term rates based on a folded unconformity.Stonewall Bank is a growing anticline that iscrossed by a stream channel cut while the bankwas subaerially exposed during the latestPleistocene sea level lowstand about 21 ka.Warping of the thalweg of the stream channelallows us to determine the Holocene uplift rateof the abrasion platform, and the slip rate of anunderlying blind reverse fault that generatesthe anticline. These rates are consistent withlong-term rates based on folding of the uncon-formity between Miocene and Pliocene strataon the continental shelf (PM unconformity).The slip rate of the blind reverse fault can beused to assess the earthquake hazard to adja-cent coastal communities.

572

Stonewall anticline: An active fold on the Oregon continental shelf

Robert S. Yeats*

LaVerne D. Kulm Chris Goldfinger

Lisa C. McNeill

GSA Bulletin; May 1998; v. 110; no. 5; p. 572–587; 11 figures.

}

*E-mail: [email protected]

Department of Geosciences, Oregon State University, Corvallis, Oregon 97331 andCollege of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon 97331

College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon 97331

Department of Geosciences, Oregon State University, Corvallis, Oregon 97331



METHODS

We use a modified raster-vector geographic in-formation system (GIS) for geologic mapping(Goldfinger et al., 1997b). The ERDAS IMAG-INE system emphasizes raster data manipulationand analysis, with a vector data model compati-ble with other systems such as ARC/INFO. Forthe Stonewall Bank area, this system integratesAMS (Alpha Marine Systems) 150 kHz sidescansonar imagery and bathymetry, digitized shelfbathymetry from a map published by the U. S.Coast & Geodetic Survey (1968), tracklines ofseismic reflection profiles, locations of bedrockand sediment samples, offshore oil-exploratorywell locations, gravity and magnetic anomalydata, seismicity, and interpretation of structuralgeologic data. Sidescan and submersible divedata were acquired in September, 1993, as part ofa National Oceanic and Atmospheric Adminis-tration (NOAA) Undersea Research Programcruise aboard the support vessel Cavalier. Duringthis cruise, AMS 150 SI (150 kHz) sidescansonar imagery was acquired at night to identifytargets for a two-person submersible,Delta, to in-vestigate the following day.

The AMS 150 system maps a swath 1 km wide,yielding an across-track pixel size of about 0.5 m.Initial processing of the sidescan data carried outduring acquisition consisted of slant-range cor-

rection and a time-varying gain correction forspherical spreading of the signal (cf. Johnson andHelferty, 1990). Position of the sidescan towfishwas calculated using the Sonar processing pack-age developed by Goldfinger et al. (1997b). Geo-logical interpretations of the sidescan imageryand the submersible dives were plotted at sea us-ing the Intergraph Microstation Computer-As-sisted Design (CAD) system and imported intothe GIS. This allowed the observer in the sub-mersible to use a geologic map based on the side-scan sonar image. The ship’s position was moni-tored by the Global Positioning System (GPS),and the submersible position with respect to theship was monitored by a transponder on the sub-mersible, allowing a scientist on the bridge to plotsuccessive positions of the submersible on the ge-ologic map during the dive. A video camerarecorded the view from the submersible duringthe entire time the submersible was in view of thesea floor; the video was supplemented by stillphotographs taken from within and outside thesubmersible. Formation attitudes were taken witha hand-held gyro compass calibrated with theship’s gyro compass before and after each dive.Rock samples were taken using a sampling clawmounted on the hull of the submersible. Thesidescan swath position and geological interpreta-tion were later imported into a GIS vector layerreferenced to the position of the ship and towfish.

Petroleum-industry multichannel seismic-re-flection profiles were used to map the underlyingbedrock structure. Time-depth conversions werebased on a synthetic seismogram constructed forthe Chevron Nautilus offshore well (Cranswickand Piper, 1992), together with sonic velocity datafrom the Unocal Grebe well at Stonewall Bank(located in Fig. 1).

TECTONIC SETTING

The Cascadia subduction zone, including theOregon offshore region, is formed by obliquesubduction of the young, warm, heavily-sedi-mented Juan de Fuca plate beneath the NorthAmerica plate (Fig. 1) at a convergence rate of 40mm/yr (DeMets et al., 1990). East of the Juan deFuca plate, the continental slope is underlain byan accretionary wedge which is actively deform-ing by folding, thrusting, and strike-slip faulting(MacKay et al., 1992; MacKay,1995; Goldfingeret al., 1997a). At latitude 45° N, the accretionarywedge is in abrupt contact with the Siletzia ter-rane, which consists of lower to middle Eocene,highly magnetic, high-velocity oceanic basalt(Siletz River Volcanics; cf. Snavely, 1987) over-lain by Eocene and younger marine strata (Tréhuet al., 1994; 1995; Fleming, 1996). The Siletziaterrane, which underlies the continental shelf,Coast Range, and Willamette Valley (Yeats et al.,

STONEWALL ANTICLINE, OREGON CONTINENTAL SHELF

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574 Geological Society of America Bulletin, May 1998

Figure 2. Fence diagram show-ing stratigraphy of the centralOregon coast and the regionsaround the Unocal Grebe (P-0093)and Chevron Nautilus (P-0103)wells, located in Figure 1. Heavylines mark sequence penetrated bywells; other stratigraphy inferredfrom multichannel seismic linesand sea-floor samples. Onshorestratigraphy from Snavely et al.(1969). Coarse-grained clastic de-posits are shaded. CRBG—Colum-bia River Basalt Group. Basalt atdeeper stratigraphic level may beCRBG sills or older flows, such asYachats Basalt at the bottom of theNautilus well.

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1996), was accreted to North America by middleEocene time (Duncan, 1982; Wells et al., 1984).The Coast Range was upwarped and theWillamette Valley downwarped following em-placement of the middle Miocene ColumbiaRiver Basalt Group.

The actively deforming accretionary wedgeand easternmost abyssal plain are cut by left-lat-eral faults striking oblique to the wedge with sliprates of 5–7 mm/yr (Goldfinger et al., 1997a).This suggests that the accretionary wedge is de-

forming not only by east-west shortening but alsoby north-south extension and clockwise rotation(McCaffrey and Goldfinger, 1995). In contrast,deformation rates in the Oregon Coast Range andWillamette Valley are very low, even though theWillamette Valley has been struck by historicalearthquakes such as the destructive 1993 ScottsMills earthquake (ML 5.6). Geomorphic evidenceof active deformation consists of deformed ma-rine terraces on the southern and central Oregoncoast (Kelsey, 1990; McInelly and Kelsey, 1990;

Kelsey et al., 1996) and different rates of streamincision, indicating varying rates of uplift in dif-ferent parts of the central and northern CoastRange (Personius, 1995). In the last 70 years, re-peated highway leveling indicates that coastalOregon has been uplifted at variable rates, from 0at Newport to 1–2 mm/yr near the ColumbiaRiver and 2–4 mm/yr in southern coastal Oregon.This short-term uplift has been interpreted as a re-sponse to elastic strain accumulation in the inter-seismic period between great subduction-zone

YEATS ET AL.


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earthquakes (Mitchell et al., 1994; Dragert et al.,1994; Hyndman and Wang, 1995).

The Oregon continental shelf is characterizedby low relief. Its surface is in large part a broadwave abrasion platform that cuts across strata asyoung as Quaternary in the Newport syncline. Atthe maximum Wisconsinan sea-level lowstand of121 ± 5 m below present sea level (Fairbanks,1989), marine abrasion was effective at depths asmuch as 20 m deeper (see discussion below). Sealevel rose rapidly during the latest Pleistocene andHolocene, preserving subaerial landforms of theabrasion platform. Tilting, warping, and flexural-slip faulting of the lowstand platform (Goldfinger,1994) are evidence of Holocene deformation, justas tilted, warped, and faulted marine terraces areevidence of deformation during the late Pleisto-cene in southern coastal Oregon (Kelsey, 1990;McInelly and Kelsey, 1990) and elsewhere.

STRATIGRAPHIC SETTING

Our summary of the stratigraphy of the OregonCoast Range in the vicinity of Newport is based onSnavely et al. (1969; Fig. 2). Stratigraphy beneaththe continental shelf to the west (Fig. 2) is based on two oil-exploratory wells (Figs. 1 and 3; cf.Snavely et al., 1982; well logs on file at the OregonDepartment of Geology and Mineral Industries),multichannel seismic-reflection profiles from the petroleum industry and from Snavely et al.(1980b) and Snavely (1987), and seafloor samplestaken by Shell Oil Company and by one of us(Kulm). Offshore microfaunal and lithologic cor-relations are based on drill cuttings examined bythe U. S. Minerals Management Service (unpub-lished reports) and the petroleum industry. Eocenebenthic foraminiferal stages (Ulatisian, Narizian)are those of Mallory (1959), and Oligocene andMiocene stages (Refugian, Zemorrian, Saucesian,Relizian, Luisian, Mohnian) are those of Kleinpell(1938). Study of coccolith assemblages (Bukryand Snavely, 1988) shows that these foraminiferalstages are time-transgressive and document waterdepth more accurately than age; accordingly, theyare indicative of age only in a very general way.

Lower Eocene oceanic basaltic basement, theSiletz River Volcanics of Snavely and Baldwin(1948), is exposed in the Coast Range northeast ofNewport (Snavely et al., 1976). It is assumed toextend offshore based on high seismic-wave ve-locity in a wide-angle seismic-reflection profile(Tréhu et al., 1994, 1995) and on aeromagneticevidence of high magnetic susceptibility beneaththe continental shelf (U. S. Geological Survey,1970; Snavely et al., 1980a, b; Fleming, 1996).

In the Coast Range, the Siletz River Volcanicsare overlain by the middle Eocene Tyee Forma-tion, in which deep-marine turbidite sandstone isthe dominant lithology (Newport section, Fig.

2). The Tyee is overlain by and interfingers withthe middle and early late Eocene Yamhill For-mation, which is predominantly siltstone with afew interbeds of sandstone. The Yamhill con-tains benthic deep-water foraminifers of the up-permost Ulatisian and lower Narizian stages(Snavely et al., 1969, 1976, 1980a). The Yamhillis overlain by tuffaceous, deep-marine siltstoneof the latest Eocene Nestucca Formation; this lo-cally overlaps older strata to rest directly on theSiletz River Volcanics. The Nestucca containsbenthic foraminifers of the upper Narizian andlower Refugian stages. Subaerial and submarinebasalt flows and breccia interbedded with theNestucca are correlated to the Yachats Basalt.

The Nestucca Formation is overlain con-formably by tuffaceous, shallow-marine silt-stone of the Alsea Formation, with Oligoceneforaminifers from the upper Refugian and Ze-morrian stages. The Alsea Formation is overlainby a deltaic sandstone unit of late Oligocene agenamed the Yaquina Formation (Goodwin, 1972).The Yaquina is overlain by and locally interbed-ded with the Nye Mudstone, predominantly adeep-marine, massive, organic-rich mudstoneand siltstone with interbeds of concretionarydolostone. Foraminifers are correlated to theSaucesian stage, indicating an early Mioceneage. The Nye is overlain with slight angular un-conformity by the early to middle Miocene As-toria Formation, a predominantly shallow-ma-rine sandstone and dark-gray carbonaceoussiltstone (Newport member of Cooper, 1981)with Saucesian (early Miocene) foraminifers.

The Astoria Formation is overlain by basaltflows of the Columbia River Basalt Group(CRBG) forming a strike ridge on the seafloorjust west of the coastline (Fig. 3; Snavely et al.,1964). Basalts of the Columbia River BasaltGroup, occurring as shallow intrusions andbasaltic breccias in lava deltas at Yaquina Headand Seal Rock and as flows farther north at CapeFoulweather, are the youngest rocks exposed on-shore in this part of the Coast Range. Two flowunits are distal parts of flows in the ColumbiaPlateau: (1) the basalt of Depoe Bay, correlated tothe low-Mg N2Grande Ronde Basalt, and (2) theoverlying basalt of Cape Foulweather, correlatedto the Ginkgo Flows of the Frenchman SpringsMember of the Wanapum Basalt (Wells et al.,1989). These basalt flows are separated by apoorly-dated massive to thick-bedded shallow-marine arkosic sandstone (sandstone of WhaleCove, Snavely et al., 1969). Offshore, sandstoneoverlies basalt west of the entrance to YaquinaBay (Snavely et al., 1964).

The offshore stratigraphy west of Newport isbased on the Unocal Grebe P-0093–1 well atStonewall Bank (total depth 3051m) and theChevron Nautilus P-0103–1 well (total depth

3849m) west-northwest of Cape Foulweather (Fig. 1). The oldest strata penetrated by the Grebewell are the Yamhill and Nestucca formations oflate middle and late Eocene age (Ulatisian andNarizian stages), with the boundary between theseformations uncertainly placed (Fig. 2). The stratathat are probably Yamhill Formation are highlytuffaceous, whereas the Nestucca Formation ismainly claystone, lacking the tuffaceous interbedsthat are found in the Yaquina Bay section. Basalt atthe bottom of the Nautilus well, below the AlseaFormation, is correlated to the Yachats Basalt ofthe Oregon coast. The Alsea Formation in bothwells is tuffaceous claystone and siltstone, as it isonshore, and it contains foraminifers of latestEocene and Oligocene (Refugian and Zemorrian)age. The Alsea is overlain by the Nye Mudstone ofearly Miocene (Saucesian) age. The deltaic sand-stone of the Yaquina Formation and the shallow-marine sandstone of the Astoria Formation pinchout offshore toward the two wells (Goodwin,1972; Cooper, 1981). Benthic foraminifers showthat the entire Eocene to early Miocene sequencewas deposited in bathyal water depths.

Basalt of the Grande Ronde Member of the Co-lumbia River Basalt Group (Snavely et al., 1980b)was encountered in the Chevron Nautilus well,where a vertical synthetic seismogram from Cranswick and Piper (1992) indicates correlationwith a prominent reflector in a seismic profilethrough the well site (Fig. 4). This same reflectoris overlapped unconformably by Plio-Pleistocenestrata east of the Grebe well (Fig. 5). In the Nau-tilus well, the basalt is both underlain and overlainby dark gray siltstone with scattered turbiditesandstone interbeds; abundant foraminifers indi-cate a middle Miocene age (Relizian stage) andupper bathyal water depths. Relizian strata werealso dredged in the core of an anticline atStonewall Bank, but were not encountered in theGrebe well. In a syncline east of the Nautilus well,the Relizian strata are overlain by more than 600m of late Miocene strata (Mohnian stage) of sim-ilar lithology. These predominantly fine-grainedstrata have no counterparts onshore. MiddleMiocene fine-grained strata cropping out on the seafloor near the coastline were deposited in 45–90 m water depths in comparison with180–600 m water depths in the Nautilus well andnorth of Stonewall Bank, suggesting that the loca-tion of the Miocene shoreline was close to (prob-ably slightly east of) that of today.

Unconformably overlying the Miocene se-quence are claystone and siltstone with shell frag-ments and fine-grained to very fine-grained sand-stone deposited in upper bathyal water depths.Benthic foraminiferal assemblages have been cor-related to the southern California Repettian andVenturian stages of Natland (1952), considered tobe Pliocene, but these stages primarily reflect wa-




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Figure 5. East-west multichannel seismic-reflection profile across Stonewall anticlineand Unocal Grebe well. Miocene-Pliocene contact (PM unconformity) based on bottomsamples. East-dipping reflectors below 3 s two-way traveltime (TWTT) are interpreted as

Alsea-Nestucca contact; these reflectors are offset by a fault, down to the west, or by a steepfold. Top: uninterpreted; bottom: interpreted. CRBG— Columbia River Basalt Group.

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Figure 6. Geologic cross section from Stonewall anticline to the Oregon coast alongthe seismic line of Figure 5. Blind reverse fault is assumed to cut the Alsea-Nestuccacontact but not the PM unconformity. YAQ—Yaquina Formation; CRBG—basaltflow of Columbia River Basalt Group; P/M—PM unconformity. Benthic microfaunal

determinations from dart core samples: MM—middle Miocene; Plio—Pliocene;Quat.—Quaternary; ON—outer neritic water depths of deposition; UB—upper ba-thyal water depths. No vertical exaggeration.

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Figure 7. Simplified cross section of Figure 6 with calculation of displacement onblind fault as expressed by deformation of the PM unconformity. Seismic data permit

the Alsea-Nestucca contact to be deformed along the steep limb of a fold or along afault. See text for calculations. A, B, C, D, E, F discussed in text.

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ter depth and have little age significance (cf.Blake, 1991). A similar sequence cored at DSDPSite 176 in 193 m water depth, southwest of themouth of the Columbia River, consisted of Pleis-tocene greenish-gray clayey silt overlying, withangular unconformity, an indurated olive-grayshale (Kulm et al., 1973). The shale contains aPliocene benthic foraminiferal assemblage indi-cating water depths ≥500 m. However, diatomsfrom the same core were assigned by Schrader(1973) to North Pacific Diatom Zone III of thelower Pleistocene. These dating problems pre-clude an age designation more precise than post-Miocene for these strata in the Newport syncline.Like the underlying Miocene strata, this sequencehas no onshore equivalent at this latitude, but inlithology and fossils, it resembles the Rio DellFormation of the Eel River basin in northwesternCalifornia (Piper et al., 1976). Water depths deter-mined from foraminiferal assemblages in theNautilus well and at Stonewall Bank are 180–600m, shallowing up-section to 45–90 m, indicatingthat rates of sedimentation became greater thanrates of subsidence. Like the Miocene sequence,the eastward shallowing based on benthicforaminifers suggests a Pliocene coastline close tothat of the present day. However, the continentalshelf was accumulating sediments during thePliocene and early? Pleistocene, whereas sedi-ments today are bypassing the shelf and are accu-mulating in large part on the continental slope andabyssal plain. Seismic profiles show that thePliocene-Pleistocene strata onlap eastwardagainst the west-sloping PM unconformity, al-though they also may crop out on the seafloor.

Piston cores from the continental slope andobservations from Deltashow that stiff, lightgray Pleistocene clay is overlain by a thin veneer of greenish-brown Holocene mud, withthe color change occurring approximately at the base of the Holocene (Barnard and McManus, 1973).

STRUCTURE OF STONEWALL ANTICLINE AND NEWPORT SYNCLINE

West of the coastline, strata are folded in theNewport syncline and Stonewall anticline.Miocene and older strata were broadly foldedand offset by normal faults prior to depositionof the Pliocene-Pleistocene sequence (Fig. 4).The largest of these normal faults, east of thesyncline, displaces the Columbia River BasaltGroup about 500 m, down to the east. East ofthe Stonewall anticline, the Pliocene-Pleisto-cene sequence overlaps the CRBG and oldermiddle Miocene strata, indicating a gentle eastdip of Eocene to Miocene strata prior toPliocene deposition (Figs. 5, 6). Two shallowsynclines west of Stonewall Bank show evi-

dence of Miocene growth, and an abrupt west-ward increase in the thickness of post-Nes-tucca strata beneath Stonewall Bank is evi-dence for pre-Pliocene displacement on theStonewall fault, discussed further below. Theserelations require a westward change, prior toPliocene deposition, from normal faulting nearthe coast to folding and reverse faulting fartheroffshore.

Faulted and folded Miocene and older strataare overlain by Pliocene-Pleistocene strataalong an unconformity (PM unconformity),with relief of up to 20 m due to Miocene foldsthat were not completely beveled by erosion.Eastward onlap of the Pliocene-Pleistocenestrata against Miocene rocks indicates that thiserosion surface sloped west about 1°, and lo-cally as much as 5°, prior to Pliocene deposi-tion. Strata on both sides of the unconformitywere deposited in water depths greater than 90 m; however, the low relief and gentle west-ward slope of this unconformity suggest that itmay have been cut by wave abrasion.

The Pliocene-Pleistocene strata are broadlyfolded into the Newport syncline (Figs. 5, 6).Structure contours based on a grid of seismic-reflection profiles (located in Fig. 3B) show thatthis syncline plunges gently about N10°W, ap-proximately parallel to a syncline in Miocenestrata below (Fig. 4). Dips on the east flank ofthe Newport syncline are 4°–12°. These dips,together with the evidence of onlap against theMiocene, show westward tilt of the west flankof the Coast Range prior to, during, and follow-ing deposition of Pliocene-Pleistocene strata.

In contrast, the west flank of the syncline,which is also the east flank of Stonewall anti-cline, dips more steeply and is not accompaniedby onlap of strata against Stonewall Bank (Fig.5). The Stonewall anticline is no more than 25km long and trends north-northwest (Fig. 3A).A seismic-reflection profile shows that the eastflank of the Stonewall anticline dips 15–18°,and the west flank dips 25°, evidence that thisanticline is seaward-vergent. A multichannelseismic-reflection profile through the site of theGrebe well shows that 1500 ± 500 m of strata,including about 400 m of Miocene strata, havebeen eroded from the crest of the anticline.North of the Grebe well about 14 km, where themiddle-Miocene outcrop on the seafloor isbroadest, 4000 ± 1000 m of strata were removedby erosion after deposition of the Pliocene-Pleistocene sequence. Uncertainty in the amountof strata eroded from the anticlinal crest allowsthe possibility of some uplift during deposition,and the possibility that the youngest strata in theNewport syncline may have been overlain by still younger sediments, now removed byerosion.

LONG-TERM SLIP RATE ACROSS THESTONEWALL ANTICLINE

Can we retrodeform the PM unconformity towork out long-term deformation rates across theStonewall anticline? The seismic profiles andoutcrop pattern of the PM unconformity showthat the Stonewall anticline is not faulted at thesurface. We interpret the east flank of the New-port syncline as part of the broad warp of theCoast Range (Fig. 7) and attribute the west flankof the syncline, Stonewall anticline, and struc-tures farther west as folds generated by east-dip-ping blind reverse faults. Thus the folding of thePM unconformity can be used to estimate thelong-term slip rate on the blind fault beneathStonewall anticline. Because we are not sure ofthe age of the post-Miocene strata other than be-ing of “Pliocene–Pleistocene” age, the long-termslip rate on the blind fault is uncertain. The foldedstrata could be entirely Pliocene, or their agecould extend well into the Pleistocene. We as-sume an age range from 4 Ma, or early Pliocene(late Repetto of southern California), to 1 Ma, orearly Pleistocene

The anticline began to form prior to develop-ment of the PM unconformity, as shown by pre-Pliocene dips in Figures 5 and 6. However, theabsence of onlap of Pliocene strata against thePM unconformity indicates that the unconfor-mity had no relief across the anticline. Further-more, multichannel seismic profiles show that thelower one-third to one-half of the Pliocene–Pleis-tocene section does not thin, and actually thick-ens locally toward the anticline from the east. Weinterpret these relations as evidence that growthof the Stonewall anticline was renewed after one-third to one-half of the Pliocene-Pleistocenestrata had been deposited. Based on our agerange of 1 to 4 Ma for the Pliocene-Pleistocenesequence, the more recent folding probably be-gan about 2 to 3 Ma.

To work out a slip rate on the blind fault, wefirst retrodeform the PM unconformity to a pla-nar surface, sloping about 1° west to accommo-date the bathymetry of Pliocene strata above the unconformity (Fig. 7). Point A on thisunconformity prior to folding (Fig. 7) would bedisplaced by folding to point B relative to thecenter of the Newport syncline, with horizontalshortening of about 400 m (A–C) and uplift ofabout 1300 m (B–C).

Nearly all of the horizontal shortening takesplace across Stonewall anticline. To determinethe dip of the blind reverse fault generating theStonewall anticline, we use the horizontal short-ening, 400 m (E–F), and the difference in alti-tude, 1000 m, of the PM unconformity betweenthe eroded crest of the anticline and the preservedtrough of the first syncline to the west (D–E, Fig.

YEATS ET AL.




7). DE divided by EF is the tangent of the dip an-gle of the blind fault. If we allow for uncertaintiesin measurement and in two-way traveltime ve-locities, the dip is 65°–70°. The east limb ofStonewall anticline is the backlimb of this fault,suggesting that fault dip decreases with increas-ing depth. Below the depth of all folding, a faultproducing 400 m of horizontal separation and1000 m of vertical separation would have a dipseparation of about 1075 m (D–F, Fig. 7).

If we assume an age of initiation of folding ofthe PM unconformity of 2–3 Ma (Pliocene), dur-ing the time of deposition, the long-term slip rateon this blind fault would be 0.4–0.6 mm/yr. Theuncertainties in the age of the onset of foldingoutweigh the errors in dip of the fault due to un-certainties in bed-length shortening and in sepa-ration normal to east-dipping bedding on bothflanks of the Stonewall anticline. The long-termslip-rate estimate assuming initiation of foldingduring Pliocene–Pleistocene deposition is lowerthan our Holocene slip-rate, discussed below, im-plying that if folding occurred during sedimenta-tion, it took place at a much slower rate than if allfolding post-dated sedimentation. The seismicprofiles show that the entire Pliocene–Pleisto-cene sequence is truncated by the abrasion plat-form, even the upper part, where there is evi-dence of minor thinning toward Stonewallanticline. If we assume that all displacement onthe blind fault occurred after the end of sedimen-tation, estimated at 1 Ma, then the slip rate on thefault would be about 1.0–1.1 mm/yr.

The difference in amplitude between the fore-limb and backlimb of Stonewall anticline may bedue to the presence of another broad synclinewith Pliocene–Pleistocene strata on the upper-most continental slope to the west (Fig. 5). Wesuspect that there is another blind reverse faultwest of the two small synclines at the west end ofthe cross section shown in Figure 6. We do notwork out the slip rate of the fault generating thisstructure because we are unable to compare long-term deformation rates with Holocene rates, aswe can at Stonewall Bank.

Another prominent reflector, probably theAlsea-Nestucca contact, dips 35°E near the Grebewell and about 15°E on the west flank of the sur-face anticline (Figs. 5, 6). This reflector has a ver-tical separation of about 2000 m at the anticlineacross a fault downthrown to the west, althoughone could also interpret the structure as an anti-cline with a steep west flank (Figs. 6 and 7). We as-sume that this fault is the same blind reverse faultthat has deformed the PM unconformity at theStonewall Bank anticline. The vertical separationof 2000 m is about twice that of the amplitude ofthe folded PM unconformity between the anticlineand the syncline to the west (D–E, Figure 7) due tofaulting and folding prior to Pliocene deposition.

The age of initiation of pre-Pliocene deformationof Stonewall anticline is too poorly constrained todetermine a slip rate based on the offset of theAlsea-Nestucca contact.

The similar amplitudes of the large synclineseast and west of Stonewall anticline raise the pos-sibility that the blind reverse faults generating thefolds could flatten into a low-dipping thrust infine- to medium-grained strata below the Alsea-Nestucca contact but above the top of Siletz RiverVolcanics. This low-dipping decollement could“collect” the blind thrusts generating all the west-vergent folds involving Pliocene–Pleistocenestrata. Considering onshore stratigraphic thick-nesses of Nestucca and older strata overlying theSiletz River Volcanics, this thrust should lie 7–10km below the surface beneath the Newport syn-cline. Yet the Siletz River Volcanics are at the sur-face in the Coast Range (Fig. 1), requiring a high-angle fault or a change to a west dip of the basaldecollement between the syncline and the coast.There is no evidence for either a high-angle faultor a decollement at the top of the Siletz River Vol-canics onshore, arguing against the decollementinterpretation.

HOLOCENE SLIP RATE ACROSSSTONEWALL ANTICLINE

The wave abrasion platform truncating thefolds of the central Oregon continental shelf issmooth enough (Figs. 3, 4, 5, 6) that the plat-form may be used as a datum for measuring tec-tonic deformation on the continental shelf. Be-fore using the abrasion platform as a datum formeasuring deformation rates, it is necessary todetermine (1) when sea level had risen enoughto remove any given part of the west-slopingabrasion platform from the erosive effects ofwave action, and (2) the water depth at whicherosive effects of waves effectively cease.

The global eustatic rise in sea level followingthe last glaciation has been documented in detailfrom coral reefs off the coast of Barbados in theCaribbean, using coral species restricted to the up-per five meters of water. Additional data for theHolocene are from Morley Island off the westcoast of Australia and the Huon Peninsula ofPapua New Guinea (cf. Peltier, 1994). Cores fromthe Barbados reefs were dated by 14C (Fairbanks,1989), and the 14C time scale in the time range10–30 ka was calibrated using high-precision U-Th ages of Barbados corals obtained by isotope-dilution mass spectrometry (Bard et al., 1990;1993). At Barbados, sea level rose by about 121 m(Fairbanks, 1989), although Peltier (1994) esti-mated that the net rise worldwide due to a globalincrease in the amount of ocean water was only105.2 m, with variations at different localities dueto the geoid effect.

The coral dating indicates that sea level rose20 m (a rate of about 3–4 mm/yr) from the lastglacial maximum (LGM) about 21 ka to about14.5 ka. Sea level then rose 24 m in less than1000 years (at a rate of more than 24 mm/yr), anevent called melt-water pulse IA (mwp-IA) byFairbanks (1989; Fig. 8, inset). Blanchon andShaw (1995) argued that sea-level rise duringmwp-IA could have been faster than 45 mm/yr,although they estimated that this high rate ap-plied to no more than about 13.5 m of the 24 m ofrise estimated by Fairbanks (1989). Dating ofcorals from Tahiti has confirmed that mwp-IAended about 13.8 ka (Bard et al., 1996). The av-erage rate of sea-level rise from the beginning ofmwp-IA at 14.5 ka to 8–6 ka, when sea level wasabout 10–15 m below that of the present day, was9–13 mm/yr. The rise in sea level during mwp-IAwas so rapid that small coastal landforms such asstream channels, sand bars, and sea cliffs couldhave been preserved on the continental shelf. Animplication of this rapid sea-level rise is that thetime at which the abrasion platform became toodeeply submerged to be further eroded (therebystarting the clock for using the abrasion platformfor deformation rates) is little affected by differ-ences in estimates of the maximum water depthswhere marine abrasion is effective.

What is the maximum water depth of effectivemarine abrasion? Bradley (1958) inferred fromthe distribution of angular and rounded pyroxenegrains off California that the maximum depth ofabrasion of wave-cut platforms was 30 feet(about 10 m) on the continental shelf. The sea-ward limit of extreme surf-related effects isknown as the “depth of closure” (Hallermeier,1981), which is the limit of the zone of extremebottom changes for beaches containing largelyquartz sand. Birkemeier (1985), using the workof Hallermeier and actual beach-profile changesat the Field Research Facility in Duck, NorthCarolina, defined the depth of closure (hc) interms of near-shore wave height (He) that is ex-ceeded only 12 hr/yr as the simple proportional-ity hc = 1.57He. An analysis of the measurementsof extreme wave heights (i.e., estimation oflargest wave height expected in some time inter-val due to a rare storm event) during the past 23years along the Oregon–Washington coast(Tillotson, 1995) reveals extreme significantwave heights ranging from 7.2 to 11.4 m for 64storms off Newport, Oregon, over a return periodof 2 to 100 years (average time period betweenstorm wave heights exceeding a threshold valueof 6 m). These extreme wave heights give a depthof closure ranging from 11.2 to 17.8 m for thecentral Oregon coast.

Based on these results, we estimate the depth ofmarine abrasion of consolidated rock, today and inthe early Holocene, as less than 20 m. Oscillatory





wave motion of long-period storm waves mayform ripples in water depths of at least 150 m (Komar et al. 1972; Kulm et al., 1975) and trans-port sand across the shelf by unidirectional cur-rents ranging up to 70 cm/sec (Smith and Hopkins,1972). However, most sand is too fine grained toabrade underlying consolidated sedimentarybedrock, as would the coarser-grained sandswhich occur on the outer shelf as relict depositsfrom the LGM lowstand (Scheidegger, et al.,1971; Kulm et al., 1975). However, bottom cur-rents laden with sand transported from CoastRange rivers and localized by topographic lows onthe shelf may have eroded the shelf in the initialstages of submarine-canyon formation.

We suggest that deformation of the abrasionplatform at the LGM lowstand at 121 m waterdepth began to be preserved at about 14.5 ka, dur-ing (or slightly prior to) mwp IA, when sea levelhad already risen about 20 m. A channel onStonewall Bank, discussed below, now at 60–70m water depth, would have been removed fromwave action when sea level was 50–60 m higherabout 11–12 ka (see inset, Figure 8). This assumes

that no differential tilting occurred from 14.5 to11–12 ka between Stonewall Bank and the LGMshoreline. However, we argue below that this sur-face has been uplifted about 15 m in the last11–12 k.y. Because sea level was rising veryrapidly when the shoreline crossed StonewallBank, adding Holocene uplift makes the age of fi-nal preservation of the platform at StonewallBank only slightly older than assuming no uplift.For purposes of slip-rate determination, we as-sume that Stonewall Bank was no longer affectedby marine abrasion after 11 ± 0.5 ka, using theBarbados sea level curve (inset, Fig. 8).

For the last 6–8 k.y., global sea level has risenfrom –10 m to the present level at a rate of about1–2 mm/yr (Fairbanks, 1989; inset, Fig. 8). Thistiming is consistent with that derived from aHolocene transgressive sequence in the GraysHarbor basin of southwest Washington, whichshows a decrease in the rate of sea-level rise from13 to about 2 mm/yr at 10–15 m below presentsea level at 8–6 ka (Peterson and Phipps, 1992).The timing is similar, although less well con-strained, in cores through the Holocene trans-

gressive sequence in Alsea Bay, east-southeast ofStonewall Bank (Peterson et al., 1984; cf.Hutchinson, 1992).

We turn now to the suitability of the abrasionplatform for measuring deformation on the con-tinental shelf. East-west profiles (Fig. 8) acrossthe Stonewall anticline at 44°30′N (StonewallBank itself) and 44°38′N, constructed from thebathymetric map along tracklines of multichan-nel seismic-reflection profiles (Fig. 3), show thatthe platform rises westward as much as 30 m be-tween the trough of Newport syncline and thecrest of Stonewall anticline, in part controlled bya strike ridge of Pliocene–Pleistocene strata eastof the axis of Stonewall anticline. The anticlinalcrest shows no evidence of upwarping along theprofile at 44°38′N, probably because it crops outat a depth shallower than the LGM lowstand, anda tectonic signal cannot be distinguished from thedifferential erosion of the strike ridge. However,in the profile at 44°44′N, the anticlinal crest cropsout at –150 m, below the LGM lowstand, and isexpressed topographically as a ridge trendingnorth-northwest (Fig. 3). Farther northwest, the

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Figure 8. East-west topographic profiles along lines of multichannel seismic sections used in structural analysis, based on U.S. Coast and Geo-detic Survey (1968). Note vertical exaggeration. X’s are at 120 m depth, the assumed maximum depth of wave action at the latest Pleistocene low-stand at Barbados. Solid lines mark fold axes. Inset: Sea-level curve for the past 19 k.y., based on coral reefs at Barbados, supplemented by addi-tional data from Morley Island, Australia, and Huon Peninsula, Papua New Guinea, from Fairbanks (1989), Bard et al. (1990, 1993), and Peltier(1994). mwp-IA marks a time of extremely rapid sea-level rise.



submarine slope is dissected by erosion (Fig. 3)due to bottom currents laden with sediment de-rived probably from the Coast Range via theYaquina and Alsea rivers. Still farther north, theanticline cannot be distinguished among severalfolds of relatively short wavelength in a multi-channel profile at 44°52′N (Tréhu et al., 1995).

In summary, the abrasion platform is not asmooth, beveled surface. It contains rocky, for-merly subaerial highlands like the strike ridgeseast of the Stonewall anticlinal axis, reachingheights of –7 m (7 m below sea level). Thesehighlands were dominated by differential sub-aerial erosion. In addition, the platform appearsto have undergone further scouring by sediment-charged bottom currents following the axis ofthe Newport syncline and dissecting the uppercontinental slope at the north end of Stonewallanticline. Finally, the abrasion platform is a com-posite of many Pleistocene lowstands. EarlierPleistocene lowstand platforms could have beendowndropped enough to be removed from sub-sequent wave abrasion, so that only the shal-lower part of the platform was abraded duringthe LGM lowstand. Accordingly, it is possible to

document deformation of the abrasion platformin a general way, but not quantitative enough todetermine deformation rates.

Fortunately, Stonewall Bank is crossed by an an-tecedent stream (Figs. 9, 10), originally subaerial,at 44.5° N. We traced the stream channel with AMS150 kHz sidescan sonar for a distance of 5.7 km,and expression on the –60 and –70 m contours onthe 1:250000-scale map (Fig. 3) permits the chan-nel to be mapped for an additional 2.3 km, for a to-tal distance of 8 km. The location and orientation ofthe channel, S65°W, suggests that it is the latestPleistocene lowstand continuation of the YaquinaRiver (Figs. 1,3; cf. Kulm, 1965). The stream chan-nel cuts across Pliocene strike ridges, and the flood-plain increases in width downstream from 275 m to550 m. There is at least one low terrace above themain channel, and side drainages enter the mainchannel from both sides. Swath bathymetry ac-quired with the AMS 150 sidescan sonar showsthat the terrace riser north of the channel is about 12m high, with a riser slope of 18° to 42°. Deltasub-mersible dives showed that the channel is nowpartly covered by Holocene mud (Fig. 9).

Bathymetry of the channel accompanying the

AMS 150 sidescan sonar shows a slope fromabout –65 m at the crest of Stonewall anticlineeastward to –72 m. Bathymetry from the1:250 000 map shows the channel expressed far-ther east by the –80 m contour, although the mapcontours shown on Figure 3 are less accurate;these show the channel as shallow as –60 m at thecrest of the anticline, 5 m shallower than the AMS150 bathymetry. In addition, the sidescan sonarimage (Fig. 9) confirms observations from sub-mersible dives that mud (dark, caused by low re-flectance) is ponding behind the crest of the anti-cline, where the channel shows higher reflectance.

The observation that the stream channel atStonewall Bank is the downstream continuationof the Yaquina River led us to construct a longitu-dinal profile from the upper reaches of theYaquina River across Stonewall Bank (Fig. 11).The profile of the Yaquina River surface fromNiem (1976) shows that the river is graded to pre-sent sea level, as are other west-flowing streams inthe Oregon Coast Range (Niem, 1976; Personius,1995). The thalweg of the river is at a water depthof nearly 4 m where the river is tidal (Kulm,1965). To this must be added an unknown thick-



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Figure 9. Sidescan sonar mosaic of eastern part of Stonewall Bank channel, 35 km offshore central Oregon. Imagery from AMS 150 kHz sonar,0.5 m resolution, and 1.0 km swath width. Channel truncates strike ridges of east-dipping Pliocene-Pleistocene strata. Low reflectivity (dark ar-eas) at east end of channel is caused by ponding of Holocene mud. Thalweg depth is –66 m at west edge of figure, and –72 m at east edge.



ness of sediments deposited during late Holocenesea level rise, as documented by geotechnicalboreholes drilled in the 1930s for the Yaquina BayBridge (logs stored at Oregon Department ofTransportation) and high-resolution seismic re-flection profiles obtained by the U.S. Army Corpsof Engineers immediately west of the coastline(Sylwester et al., 1996). For the lower UmpquaRiver, for example, these sediments are more than40 m thick (Curtiss et al., 1984), and the thalwegof the Alsea River, buried by Holocene transgres-sive sediments, is at a depth of –55m (Peterson etal., 1984). We assume in Figure 11 that the buriedthalweg of the Yaquina River is at the same depthas that of the Alsea River.

The bathymetric contours from 0 to –50 m arerelatively straight, parallel to the coastline, andshow no evidence of the channels of west-flow-ing rivers (Fig. 3). This indicates that thenearshore bathymetry is strongly affected by lateHolocene wave abrasion. Wave abrasion at shal-low depths and downcutting by sediment-charged bottom currents in the trough of theNewport syncline prevent the direct mapping of

the Yaquina River thalweg west to the thalweg ofthe Stonewall Bank channel. We estimate thatthe difference in elevation of the thalweg be-tween the crest of the anticline and trough of theNewport syncline was no more than 15 m priorto Holocene erosion.

West of Stonewall anticline, the topographyslopes westward to –72 m in the AMS 150 ba-thymetry. We estimate the depth to be about –75 m and no more than –78 m in the synclinaltrough to the west, based on the depth recorderaccompanying the east-west multichannel seis-mic line crossing Stonewall Bank and on the1:250 000 bathymetric map. Although thedepth of the channel in the synclinal trough tothe west was not mapped directly, the differ-ence of 10–13 m between the thalweg on theanticlinal crest and the estimated depth in thesyncline to the west may be used as the verticalcomponent of Holocene deformation on theStonewall blind reverse fault. Adding the tidalgradient of Coast Range rivers of 0.1–0.2m/km (Personius, 1995) would reduce this fig-ure by less than one meter.

If the 10–13 m difference were due to a blindreverse fault dipping 65–70°E, with the fault dipbased on retrodeforming the PM unconformity,the slip rate would be 0.9–1.3 mm/yr. The rangein slip rate values is due to uncertainty in the time(11–12 ka) and depth that wave action ceasedplaning off the abrasion platform and in uncer-tainty in the depth of the thalweg in the synclineto the west.

Why is the estimated depth of the thalweg ofthe stream shallower west of the Stonewall anti-cline than it is in the Newport syncline? As dis-cussed above for the PM unconformity, it is prob-ably related to another east-dipping blind-reversefault west of Stonewall Bank, west of which is asyncline on the upper continental slope containinga thickness of Pliocene–Pleistocene strata almostas great as the thickness in the Newport syncline(Fig. 5). We did not determine the shortening rateof the PM unconformity across these structuresbecause we cannot compare this rate with a de-formed Holocene feature, such as a stream chan-nel, as we can at Stonewall Bank.

In summary, displacement rates of the blind

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5000 1000Meters

SYNCLIN

E

ANTICLINE

CHAN

NEL

NADIR

STRIKERIDGES

NADIR

JOINTS

Figure 10. Sidescan sonar mosaic of western Stonewall Bank showing Stonewall Bank channel, Pliocene strike ridges, and a set of jointsstriking west-northwest, parallel to Daisy Bank fault. Thalweg depth is –72 m at west edge of profile, –66 m at east edge. First syncline west ofStonewall anticline is imaged at left; Stonewall anticlinal axis is close to the turn in the channel profile.



reverse fault deforming the PM unconformity are0.4–0.6 mm/yr if folding began during sedimen-tation and 1.0–1.1 mm/yr if folding began aftersedimentation. Displacement rates on the faultdeforming the thalweg of the stream channelcrossing Stonewall Bank are 0.9–1.3 mm/yr, ifthe fault dips 65°–70° and the channel is 11–12ka in age. Considering the uncertainties in eachestimate, we conclude that the long-term andshort-term slip rates are similar.

DISCUSSION

Origin of the Stonewall Anticline

The Stonewall anticline, considered here ascontrolled by a blind fault with a slip rate close to1 mm/yr, is only 25 km long; it is not found onmultichannel seismic lines at latitude 44°22′ Nand 44°52′ N. Its low length-to-width ratio (aspectratio) and the steep dip of the blind fault based ondeformation of the PM unconformity suggest thatthis is not part of a thin-skinned fold-thrust belt.The fold is east of the probable termination of theDaisy Bank left-lateral fault (Fig. 1), with a sliprate at the deformation front of 5.7 ± 2.0 mm/yr,and southeast of the known termination of the

Wecoma fault, with a left-slip rate, also measuredat the deformation front, of 8.5 ± 2.0 mm/yr(Goldfinger et al., 1997c).

The continental slope with the left-slip faults ischaracterized by folds of relatively short wave-length, in contrast to the long-wavelength folds ofthe continental shelf (Fig. 1). The long-wavelengthfolds of the shelf may reflect the underlying rigidbasement of Siletz River Volcanics, in contrast tothe continental slope, where this basement is ab-sent (Tréhu et al., 1994; 1995; Fleming, 1996), al-though the syncline on the upper continental slopewest of Stonewall Bank (Fig. 5) has approximatelythe same wavelength as the Newport syncline. Thewestern boundary of Siletzia should be close to theStonewall anticline, based on criteria for SiletzRiver basement outlined above (Fleming, 1996).The change in structural style may be expressed inthe short-wavelength folds immediately west ofthe Stonewall anticline (Fig. 1).

AMS 150 sidescan sonar at Stonewall Bankshows a pronounced joint pattern strikingN65°–70°W, approximately parallel to theDaisy Bank and Wecoma faults, but oblique tothe north-northwest strike of bedding (Fig. 10).However, there is no evidence of strike-slipoffset of strike ridges along any of these joint

surfaces. The joints could be a reflection ofsimple shear distributed across the entireStonewall structure, possibly because so muchof the underlying section consists of fine-grained strata.

Seismic Hazard to Coastal Communities

The Stonewall anticline is about 30 km west-southwest of the City of Newport. If the entirefault (25 km long × 20 km downdip width) rup-tured in an earthquake with average slip of 1m,regression curves of Wells and Coppersmith(1994) suggest that an earthquake of Mw = 6.8± 0.25 would be generated. An average recur-rence interval of about 1000 yr is based on slipof 1 m divided by the slip rate of 1 mm/yr, withat least a 50 percent uncertainty in slip rate andslip per event. Using attenuation relations forrock sites affected by crustal earthquakes inwestern North America (Geomatrix Consult-ants, 1995), an earthquake of this magnitudewould produce peak ground accelerations up to0.15–0.2gat 30 km source-to-site distance.This hazard assessment assumes that all slip isreleased by earthquakes, although there is no paleoseismological evidence to indicate



coastline

YAQUINABAY

LATE HOLOCENEWAVE ABRASION

base of channel

(Kulm, 1965)

DOWNCUTTING BYSEDIMENT CHARGEDBOTTOM CURRENTS

NEWPORTSYNCLINE

STONEWALLBANK

AMS 150 NOAA

15m

100

KM

SEA LEVEL

50

100

150

50

(vertical scale in meters)

? ?THALWEG OF

HOLOCENE FILL INALSEA BAY

(PETERSON ET AL., 1984)

ORIGIN

AL THALW

EG

YAQUINA RIVER (NIEM, 1976)

Figure 11. Longitudinal profile of Yaquina River from Niem (1976), extended west to thalweg of channel crossing Stonewall Bank. Original thal-weg offshore obscured by subsequent wave abrasion and scouring by bottom currents. Thalweg in tidal reaches of Yaquina River not mapped be-cause thickness of sediments deposited during Holocene transgression is not known; thalweg at base of Holocene transgressive sediments in AlseaBay (Peterson et al., 1984) shown for comparison. AMS—Alpha Marine Systems; NOAA—National Oceanic and Atmospheric Administration.



whether all slip is released seismically orwhether part (or all) is released aseismically.However, if Stonewall anticline overlies thickEocene basaltic basement of the Siletz RiverVolcanics (Tréhu et al., 1994; 1995; Snavely et al., 1980b; Fleming, 1996), it would seemlikely that this basement could store elasticstrain energy and generate earthquakes.

THE CONTINENTAL SHELF ABRASIONPLATFORM AS A STRAIN MARKER

The back-tilting of the thalweg of the an-tecedent stream channel at Stonewall Bank to-ward its onshore continuation, the YaquinaRiver, is evidence that the thalweg of thisstream channel can be used as a strain marker.To a lesser extent, the erosion surface of thecontinental shelf itself can also be used as anindication of Holocene deformation, but onlyqualitatively unless differential subaerial ero-sion, Holocene scouring by sediment-chargedbottom currents, and earlier Pleistocene low-stand wave abrasion platforms that have sub-sided to depths greater than the LGM sea-levellowstand can be taken into consideration.Downwarping may remove earlier Pleistoceneabrasion platforms from the effects of latestPleistocene marine erosion, as appears to be thecase for the erosional truncation of the synclinewest of Stonewall Bank (Fig. 5).

The age of erosion is dependent on the tim-ing and rate of eustatic sea level rise on thecontinental shelf, which is known relativelywell. Stonewall Bank and the Newport syn-cline are two structures where the abrasionplatform is deformed in the same structuralsense as underlying Pliocene–Pleistocenestrata. Mapping of the abrasion platformshould allow a very generalized structure-con-tour map to be prepared of much of the conti-nental shelf, recognizing that the preservationof the platform began at 14.5 ka at the positionof the LGM shoreline, but later at shallowerdepths (cf. inset, Fig. 8). Structures mapped inthis way could then be compared with tectonicgeomorphology of folds and thrusts fartherwest on the continental slope, below the effectsof subaerial erosion.

Lowstand wave abrasion platforms andshoreline angles can be used to map late Qua-ternary deformation, just as highstand featurescan. These platforms are commonly muchmore widespread than highstand features, oc-cupying most of the continental shelf where ithas been eroded, although they may be moredifficult to date. Lowstand platforms are notaffected by subaerial erosion after Holocenesea-level rise, although they may be furthereroded in the Holocene by sediment-charged

bottom currents that could eventually carvesubmarine canyons. The potential of mappinglowstand features may be even greater becauseof the episodic nature of sea-level rise, includ-ing the near-catastrophic mwp-IA pulse at 14.5ka. Well-dated pulses in sea-level rise permitthe measurement of deformation rates after theremoval from wave action of the abrasion plat-form and the thalwegs of rivers crossing it.

We have determined the slip rate on a blindreverse fault beneath the continental shelf basedon deformation of a horizon several millionyears old and a formerly-subaerial Holoceneriver channel now submerged in 65 m of water.The new techniques of determining slip rates atsea have widespread application on continentalshelves around the world, indicating that thecontinental shelves themselves may serve as amonitor of crustal deformation.

ACKNOWLEDGMENTS

We thank the crew of the research ship Cavalier, Deltasubmersible pilots Rich Slater,Dave Slater, and Chris Ijames, other membersof the scientific party (Gary Huftile, Cheryl Hummon, and Craig Schneider), and Kevin Redman, and David Wilson of Williamson and Associates. Williamson and Associatesprocessed the bathymetry of the channel ac-companying the 150 kHz sidescan sonar sur-vey. The U. S. Minerals Management Serviceat Camarillo, California, provided offshore datafor the study, including unpublished biostrati-graphic and paleobathymetric studies by S. Drewery, K. Piper, and H. Cousminer. Addi-tional microfaunal interpretations were pro-vided by the petroleum industry. Data for thetwo offshore wells were obtained from DennisOlmstead of the Oregon Department of Geol-ogy and Mineral Industries. Sigmund Snelsoncontributed sparker profiles and dart-core sam-ple descriptions and ages along most of theheavy lines shown in Figure 3B. Discussionswith Peter Clark and Paul Komar were usefulin determining the timing of eustatic sea-levelrise and the cessation of marine abrasion on thecontinental shelf, and Alan Niem advised us onstratigraphy of the Coast Range and continentalshelf. Ivan Wong reviewed our seismic hazardassessment of the reverse fault underlying theStonewall anticline. Reviews of the manuscriptby Alan Nelson, Tom Rockwell, and BruceSchell were very much appreciated. The studywas supported by the NOAA Undersea Re-search Program of the West Coast National Un-dersea Research Center at the University ofAlaska through grant UAF-92–0061 and UAF93–0035, by U.S. Geological Survey NationalEarthquake Hazards Reduction Program

awards 14–08–0001-G1800, 1434–93-G-2319,and 1434–93-G-2489, and by National ScienceFoundation grant EAR-96–28576 through theActive Tectonics Initiative.

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Geological Society of America Bulletincascadiageo.org/documentation/literature/cascadia... · rection and a time-varying gain correction for spherical spreading of the signal (cf