Tectonics and paleogeography of a post-accretionary forearc basin, Coos Bay area, SW Oregon, USA

This three-day field trip explores the Cascadia Forearc basin with the objective of comparing the stratigraphy and tectonics of the Coos Bay area with the regional tectonic history of Siletzia (Wells et al., 2014). The Wells et al. Geosphere paper did not include any data from our study area, so this field excursion provides new information for comparison with the most recent regional studies. The field trip area in the Oregon State Parks of Cape Arago offers spectacular outcrops of the Cenozoic history of the Coos Bay area in accessible cliffs and coves along a 16 km (10 mi.) coastal transect. Since the late 1800s, geoscientists have described these rocks and interpreted depositional settings ranging from deep-water gravity-flow facies to deltaic shoreface and coastal plain coal measures. The new generation of data is being assembled by the 23 geologists listed in the acknowledgments who are participating in the Coaledo Project that began in 2017. This research project continues as an informal eclectic assembly of collaborative researchers and students integrating a diverse spectrum of inquiry.

Each field-trip participant should have a state highway map for following our trip from Portland along Interstate 5 south to Drain on Highway 38, through the southern Coast Range to Reedsport, then south on Highway 101 to Coos Bay. The return trip will be up the coast Highway 101 to Lincoln City then along Highway 18 across the northern Coast Range and Tualatin Valley to Portland (Fig. 1).

Our tour of the Cascadia forearc will focus on the Coos Bay area, an “onlapping” sedimentary history following the 50–49 Ma accretion of Siletzia. Siletzia is a 56–49 Ma oceanic basaltic plateau that accreted to North America during the early Eocene. Wells et al. (2014) and Camp and Wells (2021) summarized the Paleogene history of Siletzia and McNeill et al. (2000) summarized the Neogene history of offshore western Oregon that sets the stage for comparison with the Coos Bay Basin history.

Our field excursion will visit the stratigraphic section exposed along the sea cliffs from Cape Arago to Gregory Point, Coos Head, and Barview areas described in Section III that provides logistic details and outcrop discussions. This area was first mapped by Diller (1901) and later in detail by Allen and Baldwin (1944). The geologic maps of Madin et al. (1995) and Wiley et al. (2015) update the structure and stratigraphy expanding on these early efforts.

Our field stops will focus on the middle to late Eocene Beds of Sacchi Beach and Coaledo Formation members that make up much of the best outcrop exposures. South Jetty construction in 1920s led to the accumulation of Bastendorff Beach and dense vegetation cover of late Eocene to Oligocene Bastendorff Shale and Tunnel Point Sandstone so isolated outcrops are only mentioned. The two Miocene formations, the Tarheel and Empire formations, will be discussed and visited if the tides allow.

The nineteen stops of this field trip merit two days for the Coos Bay area. If only one day is available Armentrout suggests that field stops 3, 5-6, 8, 10, 11-12(?) and 17-18 be prioritized. Stops 3, 10 and 17-18 require low tides.

CAUTION: Several of the stops are intertidal areas along sea cliffs. Visitors are urged to be aware of the time of high and low tides. Caution is urged, as onshore winds can significantly alter the timing and amplitude of the tide and the size of individual waves. The stop sequence is in stratigraphic order but should be adjusted to fit the tide and weather situations at the time of the field excursion. Several stops are at sea cliff viewpoints. These areas are underlain by Pleistocene terrace sands and are often undercut. Approach all bluffs with great caution.

Overnight accommodations are available at Coos Bay and Charleston. Campgrounds are available at Sunset Bay State Park and Bastendorff County Park. Restroom facilities exist (if open) at Cape Arago, Shore Acres, Sunset Bay, and Bastendorff Beach.

Departing Portland, we leave the Portland Basin and traverse the Willamette Valley driving south on Interstate 5 (Fig. 1). The fault-bound Portland basin is filled with Miocene and younger strata (Evarts et al., 2009). These rocks include sediments from the Columbia River, flows of the Columbia River Basalt Group, glacial-flood deposits of the Missoula Floods, and between 32 and 50 cinder cones and basaltic volcanoes of the Boring Lava that form many of the conical hills in east Portland (Fig. 2).

The Willamette Valley is the eastern segment of the Cascadia forearc between the Cascade volcanic arc to the east and the uplifted Coast Range to the west. The valley floor is blanketed with gravels, sands, and silt of the Missoula Floods (19–15 k.y.), making wonderful agricultural land for grains and livestock (Fig. 3). The largest Missoula Floods inundated the Willamette Valley as far south as Eugene flooding Portland to a depth of ~90 m (300 ft) (Allison, 1935; Bretz, 1969, Waitt, 1985; Minervini et al., 2003; Waitt et al., 2019).

The uplifted Paleogene marine sedimentary rock coring the Coast Range weather into soils perfectly suited for grapes supporting many of the more than 500 Oregon wineries. Just south of Salem, Oregon’s state capital, we cross the Salem Hills where flows of the Columbia River Basalt Group (15 Ma) are weathered into bright-red laterites suggesting “tropical” conditions during part of the middle Miocene. These flood basalt flows erupted from fissures in eastern Oregon and Washington, flowed within ancestral Columbia River valleys through the Cascades and as far as the Pacific Coast where ponded lava invasively intruded downward into water-saturated sediment (Fig. 4). See Beeson et al. (1989), Wells et al. (2009), and Reidel et al. (2013) for the Columbia River Basalt story.

South of Albany and the highway exit to Corvallis, home of Oregon State University, we begin to see east of the freeway rounded hills of the Little Butte Volcanics (35–25 Ma), erupted during an early phase of the Western Cascade volcanic arc. Several of these hills are exhumed intrusions and provide quarry rock for highway construction. Off to the west is the highest peak of the Coast Range, Marys Peak (1249 m; 4098 ft). The lower slope of Marys Peak consists of the Siletz River Volcanics (56–49 Ma), part of the oceanic basalt plateau terrane of Siletzia that docked with proto-Oregon ca. 51–49 Ma. The upper portion of Marys Peak is composed of Tyee Formation sedimentary rocks intruded by a post-accretion sill.

The city of Eugene, home of the University of Oregon, is underlain by fossiliferous sandstone of the Oligocene Eugene Formation and volcaniclastic Fisher Formation of floodplain and lake deposits. At the Goshen exit just south of Eugene, an outcrop east of the freeway exposes Fisher Formation lake deposits including a re-sedimented white tuff layer and debris flow deposits that are intruded by three basalt dikes (Retallack et al., 2004). As the Coast Range began to rise, the marine waters of the Willamette Valley were displaced northward until the shoreline was west of Portland by the Miocene (Orr and Orr, 2012).

At the Drain exit for Highway 38 to Reedsport, we encounter the Tyee Formation sandstones that underlie much of the uplifted Coast Range. These middle Eocene sandstones are gravity-flow deposits sourced from the Klamath Mountains. The suturing of Siletzia against the accreted Mesozoic terranes at 51–49 Ma created a fold-and-thrust belt involving some of the pre-forearc sedimentary units (Umpqua Group) (Wells et al., 2000). The older Klamath Mountain rocks of this fold-and-thrust belt shed sediments across a broad subtropical coastal plain and into the Cascadia forearc basin progressively burying the seamount volcanic centers of Siletzia (Santra et al., 2013) These sandstones and interbedded shales core the steep hills along the 90 km (56 mi.) drive to Reedsport. There is a small post-Tyee inlier of the prodelta Elkton mudstones exposed at the summit of Highway 38 just before descending into the town of Elkton. Overlying the Elkton are deltaic facies of the Bateman Formation, a correlative of middle Eocene deltaic facies within the Coos Bay area (Ryu et al., 1996; Santra et al., 2013).

Highway 38 follows Elk Creek and the Umpqua River to Reedsport. The Umpqua River headwaters are near Crater Lake in the high Cascades. The Umpqua and Rogue rivers have maintained their pathways to the Pacific Ocean from at least earliest Miocene time. Both the Umpqua River and Elk Creek have superimposed a locally meandering pathway now in deep valleys as the Coast Range uplifted.

From Reedsport south to Coos Bay we leave Tyee Formation sandstones and enter the Coos Bay area where the middle Eocene Coaledo Formation contains coal mined into the 1940s and where timber has been the main economic driver peaking in the 1970s. Recent exploration for Coal Bed Methane discovered potential resources that could become economic if the current efforts are successful to build a liquid natural gas terminal in the Coos Bay area.

Coos Bay area strata onlap the accreted terrane of Siletzia (Fig. 5) and are the westernmost extension of the Paleogene Cascadia forearc basin (Fig. 6) that onlaps the southern margin of Siletzia. The literature on this convergent margin is diverse, with new details and creative hypotheses emerging monthly. The primary objective of this field guide is to visit the rocks of the Coos Bay area that expose a relatively complete but punctuated record of Cascadia forearc events.

The following text summarizes the events and timing detailed in Wells et al. (2014), Camp and Wells (2021), and McNeill et al. (2000) with additions based on Coos Bay area facies associations and zircon data (Armentrout et al., 2020). Details with selected references follow this summary.

Key to understanding the Cascadia Forearc history are five primary phases:

1. Siletzia Terrane Extrusion; 56–49 Ma: During the early Eocene, “Siletzia,” an oceanic basaltic plateau, was extruded on the Farallon plate and accreted to the North American plate at 51–49 Ma ending the Challis-Absaroka subduction zone. The Yellowstone hotspot, inferred to be a whole-mantle plume, is considered as the magma source for Siletzia.

2. Paleocene Cascadia Basin Initiation; 48–45 Ma: Following accretion of Siletzia, the new 49–46 Ma western Cascadia subduction zone formed an early middle Eocene accretionary prism that became the western boundary of the transitional pre-arc Cascadia basin. The subducting Farallon slab reached melting temperature depths initiating middle Eocene volcanism as early as 45 Ma in Washington and ca. 40 Ma along the Oregon segment in the Cascade volcanic arc.

3. Paleocene Cascadia Forearc Extension; 46–30 Ma: The Klamath Mountain folded-and-faulted suture zone uplifted mountainous areas from which abundant sediments were transported across a broad subtropical plain and forearc slope. These transported gravity-flow sediments buried much of the southern Siletzia seamount topography. The Coos Bay Eocene and Oligocene strata were a coeval western extension of the Tyee Basin of numerous authors. This initial Paleogene phase of Cascadia basin deposition consists of relatively conformable progradational successions of prograding coastal plain facies into a deep-water basin. Numerous unconformities occur associated with growing structures and several are interpreted as having “regional” extent.

4. Late Paleogene–Neogene Compression; <30 to >25 Ma: late Oligocene–early Miocene compressive deformation uplifted and truncated the forearc Paleogene strata and focused sedimentation into offshore depocenters with small onshore extensions. In the Coos Bay area, this deformation formed the South Slough Syncline in which the shallow marine to estuarine facies of the early to middle Miocene Tarheel formation and unconformably overlying late Miocene Empire Formation were deposited. This regional oblique-compression initiated uplift of the southern Coast Range, isolating the Coos Bay area from river transported Cascade arc sediment.

5. Increased Right/Lateral Shear; 16 Ma–Present: Progressive right/lateral shear, driven by Basin and Range extension and northward migration of the San Andreas fault system, faulted western Oregon along northwest-southeast trends and segregated upper plate areas into clockwise rotated fault-bound elements. This tectonic phase may have begun somewhat earlier but new data needs to address late Oligocene events. The Yellowstone hotspot emerged east of the young Cascade volcanic arc where crustal extension resulted in extrusion of the middle to late Miocene 16.5–12 Ma Columbia River Basalts.

The above summary is an oversimplification but captures the essentials of Cascadia forearc evolution, Paleogene forearc sedimentary basin formation, and Neogene deformation fragmenting the forearc into separate depocenters. The following text provides additional details with selected references. Three important qualifications for this section of field-trip text are as follows.

• The Paleogene geology of the southern Oregon Coast Range is referred to several “basins” based on areas of preserved strata. This includes the Umpqua Basin, Tyee Basin, Coos Bay Basin, and Sutherland Subbasin. The Umpqua Basin predates accretion of Siletzia and is clearly a distinct tectonic episode. The Tyee, Coos Bay, and Sutherland “basins” are geographically distinct areas of coeval but of differentially preserved strata of the early deposition of the pre-arc and forearc Cascadia Basin. This field guide will refer to this middle to late Eocene basin-wide depositional phase as the Cascade Basin and the successions of the Tyee, Coos Bay and Sutherland strata as “areas” rather than as basins. This evolving “basin” nomenclature evolves from discussions amongst the Coaledo Project Team, principally R.B. Dorsey, M. Darin, and J.M. Armentrout (written commun., July–August 2021).

• Paleomagnetic analyses of Cascadia forearc sedimentary and igneous rocks and modern GPS data (cited below) document a long history of right/lateral deformation resulting in several generations of clockwise rotation. That will be addressed in the subsection on Neogene deformation. The text that follows refers to paleogeographic directions in today’s orientations, not counterclockwise restoration.

• We introduce two working hypotheses that modify previous Coos Bay area and Cascadia forearc interpretations. We invite trip participants to challenge these ideas:

  1. Extension of the Paleogene Tyee-Coaledo basin boundary beyond the Fulmar fault based on analysis of microfossil and zircon data; and

  2. The initial onset of Oregon Coast Range uplift as early as <30 to >25 Ma based on zircon populations and interpreted formation ages.

Mesozoic subduction along western North America added a composite of terranes called the Saint Elias orogen to the North American plate (Fig. 5). Among the last terranes to accrete were Siletzia and Yakutat, both oceanic seamount plateaus. The Siletzia basaltic plateau terrane was extruded on the Farallon plate during 56–49 Ma and the Yakutat terrane on the Kula plate (Duncan, 1982; Wells et al., 2014; Camp and Wells, 2021) (Figs. 6 and 7). The Farallon plate portion accreted to Oregon, Washington, and Southern Vancouver Island, British Columbia, as Siletzia (51–49 Ma) (Fig. 8). The Yellowstone hotspot is considered the probable magma source for both Siletzia and Yakutat terranes (Figs. 7–9) (Camp and Wells, 2021). Figure 8 diagrams the docking of Siletzia into the Columbia Embayment and subsequent westward shift of subduction from the Challis-Absaroka subduction zone to the Cascadia subduction zone.

Siletzia docked at 51–49 Ma with the eastern margin of the Columbia Embayment bound by Mesozoic terranes of the Klamath and Blue mountains (Fig. 8 at 48–45 Ma) (Camp and Wells, 2021). The accretion of Siletzia resulted in waning of the Challis-Absaroka subduction system and initiation of the Cascadia subduction ca. 48–45 Ma (Camp and Wells, 2021). Onset of Cascadia subduction would have initiated an accretionary prism that formed the western margin of the Cascadia forearc ca. 47-45 Ma.

The deformation associated with Siletzia docking is exposed in the Umpqua fold-and-thrust belt southeast of Roseburg, Oregon (Fig. 10) (Wells et al., 2000, 2014). Major docking deformation is recorded by the 51 Ma unconformity between lower and upper Umpqua Group sediments and a later pulse at 49 Ma between the upper Umpqua Group and Tyee Formation (Figs. 10 and 11) (Wells et al., 2014). The eastern docking margin of Siletzia with North America is likely buried beneath the Cascadia volcanic arc. The western boundary is traditionally identified by gravity and magnetic data and by the northward extension of the Fulmar fault (Snavely, 1987; Clarke, 1992; Wells et al., 2014; see the discussion in the section titled “Western Boundary of Paleocene Deposition”). The northern boundary is on Vancouver Island where the Metchosin igneous complex is included as part of the accreted Siletzia terrane (Wells et al., 2014).

The Paleogene Cascadia depositional basin extended from at least the eastern margin of the Willamette Valley westward across the Coast Range to the Cascadia accretionary wedge outer-arc high (Fig. 12A) (McNeill et al., 2000). Pre-arc transitional sedimentation ca. 45 Ma began in a shelf basin as soon as the initial accretionary prism formed following initiation of the Cascadia subduction zone ca. 49-46 Ma (Camp and Wells, 2021). Forearc sedimentation in southern Oregon initiated with the onset of Cascade arc volcanism at ca. 40 Ma (Retallack et al., 2004). Regional margin-parallel-to-oblique extension accommodated thick Paleogene sedimentary deposits (ca. 45–30 Ma) (Wells et al. 2014; Camp and Wells, 2021). In the Coos Bay area, preserved middle Eocene to late Oligocene strata are more than 3400 m (11,000 ft) thick. Preserved strata in the south-central Cascadia basin (Tyee area) measure more than 7600 m (25,000 ft) of early Eocene to middle Eocene sediments (Ryu et al., 1996). This accommodation required rapid subsidence of the sedimentary basin overlying the accreted Siletzia terrane. Extension of the upper plate is demonstrated by igneous dikes with southwest to northeast trends (present day orientation) through which the Yachats Basalts and Tillamook Volcanics intruded (42–34 Ma—Tillamook stage) (Fig. 11). These dikes and associated volcanics are interpreted as sourced by the Yellowstone mantle plume as it passed beneath the overriding North American plate (including the accreted Siletzia terrane) (Camp and Wells, 2021; see Fig. 9 herein).

Basin scale schematic cross sections and maps for the Tyee and Coos Bay areas (“basins”) and paleogeographic maps for the Paleogene are illustrated in Figures 12–17. Figures 12, 13, and 14 illustrate successively more detailed cross-sections for the Cascadia forearc. Figure 13 for the Newport Basin places the “forearc” depositional boundary at the Fulmar fault inboard of the outer-margin high (Snavely et al., 1980; Snavely and Wells, 1966). In contrast both Figures 12 and 14 show the depositional boundary at the outer arc high underlain by the subduction zone accretionary prism.

The Figure 10 cross section of the Umpqua fold-and-thrust belt illustrates the deformation of Franciscan Complex and Klamath Basement that formed uplands where erosion supplied sediment transported into the developing forearc Cascadia basin. Regional geology demonstrates the pre-forearc and forearc basin was a single broad basin extending from the western margin of the Western Cascades to the outer arc high formed by the Cascadia accretionary prism and thus includes the Paleogene stratigraphy of the Coos Bay area (fig. 12A from McNeill et al., 2000) (Weaver, 1945; Dott, 1966; Snavely, 1987; Snavely and Wells, 1996; Ryu et al., 1992, 1996; Santra et al., 2013).

Late Oligocene to middle Miocene <30 to >25 Ma compressional deformation uplifted the axis of the forearc forming the southern Oregon Coast Range (Fig. 12B). This uplift eroded much of the Paleocene strata along the Coast Range axis with the younger Paleogene section preserved on the downdip western and eastern flanks and along the Coast Range crest near Elkton, Oregon. The “preserved” strata on the southwestern Coast Range flank are mapped as the Coos Bay (Basin) area and on the southeastern flank occur in the Sutherland (Subbasin) area. The southern Oregon Coast Range geologic map of Ryu et al. (1996) illustrates this preservation pattern (Fig. 15).

The Cenozoic section preserved in the Coos Bay area has traditionally been dated using molluscan and foraminiferal fossils interpreted by various paleontologists (Turner, 1938; Weaver, 1945, Bird, 1967; Baldwin, 1973; Rooth, 1974; McKeel, 1979, 1980; Snavely et al., 1981; Barron, 1981; and others). The Coaledo Project has initiated zircon dating of the sandstones. Figure 16 shows the initial results with zircon-sampled sandstones plotted on a composite stratigraphic column and with a detrital zircon chronostratigraphy for the Coos Bay (Basin) area. Michael Darin, University of Nevada–Reno, has constructed this zircon age stratigraphic framework that the Coaledo Project team is using for age correlations. The zircon-age calibration confirms most earlier biostratigraphic ages but with greater precision and facilitates chronostratigraphic correlations overcoming time-stratigraphic facies correlations.

The chronostratigraphy shows the conformable Paleogene section spanning the middle Eocene to early Oligocene with two Neogene sandstone units bound by unconformities (Fig. 16). Formation names shown here and discussed in the following text are reviewed in Baldwin (1973), Armentrout et al. (1983), Madin et al. (1995), and Wiley et al. (2015).

Figure 17 shows the Cascadia pre-forearc paleogeography illustrated by Orr and Orr (2019; from references cited there in) for the early middle Eocene Tyee-system (46–45 Ma at ca. 45 Ma), the late middle to early late Eocene forearc Coaledo-system (45–38 Ma at ca. 41 Ma), and the late Eocene–Oligocene Bastendorff–Tunnel Point system (35–30 Ma). These illustrations provide carefully crafted depositional settings based on regional biostratigraphic correlation and sedimentologic interpretations (see also Dott, 1966; Ryu et al., 1992, 1996; and Santra et al., 2013).

The Tyee “Forearc” Basin is primarily outlined by exposed Tyee Formation sediments overlying the older Umpqua Group rocks (Fig. 15). The zircon ages for the Tyee Formation suggest it was deposited between ca. 46-45 Ma, predating the onset of volcanism in the southern Oregon Cascades at ca. 40 Ma (Fig. 16 and Retallack et al., 2004). Discussions with M. Santra and R. Wells during 2021 affirm no recognized tuffs or tuffaceous sediments in Tyee strata. This would place the Tyee Formation deposition within the pre-arc transitional phase of the Paleogene Cascade Basin.

The older preserved strata of the early Eocene Umpqua Group and the overlying middle Eocene Tyee Formation are referred to as the Umpqua Basin and the Tyee Basin respectively (Figs. 15 and 17A). The younger middle Eocene to middle Oligocene section is referred to as the western Coos Bay Basin and eastern Sutherland Subbasin (Figs. 15 and 17B). The preservation of the Elkton Formation mudstones and Bateman Formation sandstones near Elkton, Oregon, represent a Coast Range axial in-liner preserving a portion of the younger Eocene Cascade Basin deposition (Fig. 15; Ryu et al., 1996; Weatherby, 1991). Coeval late middle Eocene to middle Oligocene strata most likely extended across the forearc basin to the foothills of the Cascade arc but were removed from most of the Coast Range axis by late Oligocene and younger uplift and erosion. This interpretation is essentially the same as Dott (1966) who created paleogeographic maps showing successive coastal plain facies prograding into deep basins during both Tyee and Coaledo-Bateman-Spencer phases.

Reconstruction of the western margin of the Paleogene forearc basin is based on seismic reflection profiles calibrated by offshore wells (Figs. 13–16). The Newport transect of Snavely et al. (1980; our Fig. 13) provides a model for the shelf depocenter boundary of Siletzia. The western boundary of Siletzia is identified by gravity and magnetic data and by the northward extension of the Fulmar fault (Snavely, 1987; Clarke, 1992; Wells et al., 2014). The basin margin model presented here would extend the depositional margin ~40 km west of the Fulmar fault (Fig. 18).

Seismic data from the Eel River Basin surveyed from the shelf basin across the outer arc high and down the slope across the subduction zone provides a model for the Cascadia forearc outer shelf to abyssal sea floor (Fig. 14) (Biddle and Seeley, 1983). The Biddle and Seeley (1983) seismic profile was migrated using algorithms that geometrically repositioned returned signals restoring the recorded data to the correct spatial position more precisely imaging the complex geology including the Paleogene section. The seismic data from the Newport Basin (Snavely et al., 1980) and Coos Bay Basin (Clarke et al., 1985; Clarke, 1992) was not migrated and sea-floor multiples obscured most of the Paleogene section. We have used the Eel River basin seismic transect for modeling the Coos Bay area outer shelf high as the western margin of the Paleogene “greater Tyee–Coaledo” Cascadia Basin (Fig. 18).

By analogy with the Eel River Basin geometry, we infer that the formation of the offshore Coos Bay depositional basin commenced with the first accretionary prism resulting from Cascadia subduction, estimated at ca. 47–45 Ma (Camp and Wells, 2021) (compare subduction zones between Fig. 8 at 55 Ma versus Fig. 8 at 48–45 Ma).

Early models of the Cascadia forearc used the Fulmar fault as the western boundary (Fig. 13). Snavely et al. (1980), and Snavely (1987) interpreted arkosic wacke well-cuttings of Penutian Stage age in the Union Oil Fulmar #1 and Pan American Coos Bay #1 wells (Fig. 18) to be transported northward from an unidentified southern terrane toward California. These authors used the petrographic analysis of these well cutting samples to suggest a separate terrane west of Fulmar fault (originally Fault A, Snavely et al., 1981) inferring at least 200 km (124 mi.) of right-lateral offset. The onshore Fulmar fault is a probable southwestern terrane-boundary fault with Siletzia to the north and accreted older terranes to the south (Wiley et al., 2015). Wiley et al. (2015) correlate the Pan American and Fulmar Penutian arkosic wacke well-cuttings to their Sandstone of Fivemile Point at Fivemile Point, near Bandon, 10 km (6 mi.) south of the Pan American well. M. Darin (personal commun, Sept. 2020) interpreted a zircon age for the Sandstone of Fivemile Point at maximum depositional age–youngest single grain (MDA-YSG) 52.3 ± 1.2 Ma (from data in Wiley et al., 2015) and a zircon age for the White Tail Ridge Formation of the lower Umpqua Group at MDA-YSG 48.4 ± 3.6 (data from Stern and Dumitru, 2019). Both samples correlate with the late Penutian (52–49 Ma; McDougall, 2021) and have similar composite probability density age plots of zircon ages suggesting similar if not the same zircon source terrane.

If the Sandstone of Fivemile Point and the White Tail Ridge lithofacies have the same source and are both late Penutian in age (52–49 Ma; McDougall, 2021), they are probably of the same depositional phase of early Eocene lower Umpqua Group sourced from the Klamath Mountain provenance, transported north through fluvial-deltaic systems then into deep-water by gravity-flow processes (see depositional model of Santra et al., 2013, their figure 16–Stage 1). These deposits were transported northwest, seaward of the approaching Siletzia plateau.

The interpreted direction of Siletzia’s movement in southwestern Oregon was to the southeast (today’s orientation) which could account for the right-lateral offset within the Fulmar fault deformation zone at Fivemile Point and possibly along the western edge of Siletzia (Wiley et al., 2015; Snavely et al., 1980; Snavely and Wells, 1996). As mentioned above, the western edge of Siletzia is identified by gravity and magnetic data, which is also used to define the northward offshore trend of the Fulmar fault interpreted to be active in the Eocene, but the Paleogene section in unmigrated offshore seismic data is “masked” by sea floor multiples (Snavely, 1987; Clarke, 1992; Wells et al., 1998). Right-lateral drag along the fault could account for the vertical to overturned folding of the Sandstone of Fivemile Point at Fivemile Point (Wiley et al., 2015). The Fulmar fault may be a “tear fault” along the western margin of Siletzia, complimentary to the Wildlife Safari fault along the eastern margin of Siletzia (Wells et al., 2000). This suggests that the early Eocene strata west of the Fulmar fault are part of the pre-docking depositional system of the lower Umpqua Group White Tail Mountain facies and not a northward translated terrane. This interpretation must be classed as a working hypothesis.

Benthonic foraminiferal faunas of Early Ulatisian age (49–46 Ma; McDougall, 2021) are absent in the Pan American P-0112 well and the Ulatisian (49–43 Ma, McDougall, 2021) is entirely absent in the Union Fulmar P0130 well (McKeel, 1979, 1980). This unconformity is correlative with Wells et al. (2014) timing of Siletzia accretion. In the Pan American well the unconformity overlies outer neritic foraminiferal biofacies of early Ulatisian age (49–46 Ma) and is overlain by bathyal late Ulatisian-Narizian (46–40 Ma) bathyal biofacies suggesting a deepening above the unconformity (biofacies from McKeel, 1980; ages from McDougall, 2021). Loading of the forearc margin by accretion of the basaltic Siletzia plateau and onset of initial subduction might result in such an unconformity and paleowater-depth deepening. The Ulatisian unconformity in the Fulmar well has bathyal biofacies below and above and provides no indication of paleowater depth change (McKeel, 1980).

The offshore Fulmar and Pan American wells contain strata of upper Ulatisian, Narizian, and Refugian Stage foraminifera correlative with the onshore Coos Bay area section (McKeel, 1979, 1980; Snavely et al., 1981) suggesting that the Coos Bay area, and thus the larger concept of the “Tyee–Coos Bay” Cascadia Basin, extends toward if not to the initial uplifted forearc accretionary prism of the Cascadia subduction zone as depicted on Figure 18. This encourages that maps depicting the paleo–Tyee “Basin” area should encompass the more regional occurrence of coeval facies rather than only the mapped outcrop region of the Tyee Formation (contrast Fig. 15 with Fig. 18). This interpretation is similar to that of McNeill et al. (2000) who extended the Newport area Paleogene basin to the outer-arc high formed above the accretionary prism (Fig. 12B).

The “Tyee-Coos Bay” Cascadia Basin depositional system shows generally northward progradation from the Klamath Mountains across a broad subtropical (Hopkins, 1967) coastal plain into a subsiding forearc basin onlapping onto the seamount topography of Siletzia (Fig. 17A; ca. 45 Ma) (see also Santra et al., 2013). The Umpqua Arch (Fig. 15) is interpreted as a Siletzia-terrane volcanic high (Ryu et al., 1996), a basement high (Santra et al., 2013), or as a foreland bulge (R.J. Dorsey, personal commun., Dec. 2020 and Jan. 2021), separating a southeastern Umpqua Group depocenter from the more distal northern area. The northeast-to-southwest oriented Umpqua Arch projects offshore toward the Pan American well P-0112 and may be related to the unconformity between Penutian and Ulatisian Stage strata (see previous section discussion of the “Western Boundary of Paleogene Deposition”).

The Umpqua Arch “basement” high was prograded across by distal uppermost Umpqua Group and lower Tyee Formation gravity-flow sandstones as the depocenter moved farther north (see Santra et al., 2013; Figs. 16 and 18). Integration of regional geology with seismic and well data lead to the evolution of forearc basin depocenters being referred to as the Umpqua Basin, Tyee Basin and Coos Bay Basin, all being successive tectono-stratigraphic phases of Cascadia Paleogene forearc sedimentation. Numerous unconformities punctuate this Paleogene stratigraphy, some relatively regional and others well developed on local structural highs then becoming conformable in adjacent depocenters (Baldwin, 1973; Ryu et al., 1992).

The northward progression of a generally west-southwest to east-northeast coastline is envisioned for each prograding depositional phase; the Umpqua phase, Tyee phase including Beds of Sacchi Beach and lower Coaledo, and Elkton/Bateman phase that is likely coeval with Middle and Upper Coaledo to Tunnel Point strata at Cape Arago. This paleogeographic trend, described with modern compass coordinates, emerges from Dott (1966) and Santra et al. (2013) study of the Paleogene forearc. Late Oligocene into the early Miocene uplift of the Coast Range has differentially eroded the crest of the Coast Range down into middle Eocene Tyee facies with the younger rocks exposed predominantly along the east flank in the Willamette Valley (Sutherland subbasin of Fig. 15) and western flank along the modern coast especially at Cape Arago (Coos Bay Basin of Fig. 15), and in a Coast Range crestal inlier of Elkton and Bateman strata near Elkton, Oregon (see axial area of Tyee Basin centered over Umpqua Arch, Fig. 15) (Weatherby, 1991; Ryu et al., 1996).

The late Oligocene to early Miocene of the Cascadia forearc basin is an interval of significant regional change. The Paleogene was dominated in regional margin-parallel to oblique extension with depositional systems accumulating across the entire forearc. The Neogene is characterized by compressional and rotational tectonics with deposition within sub-regional faulted compartments along the flanks of the uplifted Coast Range. This is a period marked by offshore plate reorganization, a relatively quiet volcanic interval and onset of significant right/lateral faulting throughout the forearc from the accreted prism to the western Cascade volcanic arc (Goldfinger et al., 1992, 1997; McNeill et al., 2000; Wells et al., 2014; Wells and McCaffrey, 2013; Camp and Wells, 2021). Within the fault-compartmented onshore basins and the offshore depocenters multiple discontinuities reflect ongoing episodes of deformation and associated Cascadia volcanic arc eruption (McNeill et al., 2000; Priest, 1990; Camp and Wells., 2021) (see Fig. 9 color-coded eruptive phases).

In the Coos Bay area, the <30 to >25 Ma compressive and right/lateral faulting phase is marked by southwest-vergent thrust faulting and folding that uplifted and truncated the Paleogene section (Fig. 19). The northwest-to-southeast South Slough Syncline became the center of Neogene deposition. The early to middle Miocene Tarheel formation and the unconformably overlying late Miocene Empire Formation (Figs. 19 and 20) are interpreted as shallow marine to estuarine deposition in the northwest to southeast embayment surrounding the South Slough Syncline (Fig. 20). Ongoing Quaternary deformation again folded the Empire Formation along the axis of the syncline and has uplifted and truncated the entire Paleogene and Neogene section (Fig. 19). The magnitude of Quaternary uplift is recorded in the succession of marine terraces exposed along the coastal hills of Cape Arago and south to Cape Blanco (see Figs. 21 and 22) (McInelly and Kelsey, 1990; Madin et al., 1995; Wiley et al., 2015).

At the scale of the entire Pacific Northwest, the onset of Basin-and-Range extension, northward compression by the San Andreas fault system, driving Siletzia northward resulted in differential clockwise rotation of the Cascadia forearc (see Fig. 8 at 35–15 Ma and at 15 Ma–Present). “Basin-and-Range extension began at 17–16 Ma (Colgan and Henry, 2009) when torsional stress was fully imposed on the continental interior due to plate-boundary tectonics (Dickinson, 1997)” (from Camp and Wells, 2021, p. 7). Evidence for rotations is documented in landmark papers by Simpson and Cox (1977), Wells (1989), Wells and McCaffrey (2013), and others, demonstrating a long Cenozoic history of differential rotation that increases in magnitude from east to west reflecting the increased “compressional bulge” driven by the northwest push of Siletzia against the backstop of British Columbia accreted coastal terranes (Wells and Simpson, 2001).

The northward migration of the 400 km long Siletzia terrane due to Basin-Range extension and Pacific–North American dextral shear is calculated at up to 9 mm/yr. (Wells et al., 1998). This amounts to ~9 km/m.y. and ~150 km of compressive shortening of Siletzia since 17 Ma onset of Basin-Range and San Andreas fault system regional deformation.

Maps of the river basins of southwestern Oregon show that the Coos River and Coquille River drainage basins are restricted to the west side of the Coast Range in contrast with the Umpqua and Rogue rivers which cross the Coast Range (Fig. 23). The Umpqua River heads near Crater Lake and drains areas of the Cascade volcanic arc. The Rogue River rises in the Cascades and flows through parts of the Klamath Mountains metaplutonic accreted terranes. Segments of the Umpqua River are meanders superimposed within deep canyons across the Coast Range. These drainage patterns raised the question of “if and when” the Coos Bay area may have become “sedimentologically isolated” from the Cascadia magmatic arc as the southern Coast Range began to be uplifted due to compression and underplating.

Uplift of the northern Coast Range is occasionally interpreted as post–Columbia River Basalt Group flows that reached the coast at ca. 16.5–14.5 Ma at Seal Rock, Depo Bay, and Cape Foulweather, Oregon (Beeson et al., 1989; Wells et al., 2009; Reidel et al., 2013). The transition in the Cascadia forearc from north-south extension to southwest-northeast compression about <30 to >25 Ma might be another time for uplift. Alternatively, underplating might be the driver for uplift (Ducea et al., 2009; see their comment on Cascadia, p. 19; also see Menant et al., 2020).

Figure 24 displays the composite probability density age plots for all initial Cenozoic zircon grains from the Coos Bay Basin (Armentrout et al., 2020). Michael Darin, University of Nevada–Reno, processed sandstone samples and interpreted zircon data from the Arizona LaserChron Center at the University of Arizona (M. Darin, written commun., Dec. 2020, May 2021, and July 2021). The youngest zircons in Eocene and Oligocene units get progressively younger up section, which is an expected result in sedimentary basins that receive a continuous supply of new zircon from an adjacent and active magmatic arc. However, the detrital zircon samples from Miocene formations reveal two surprising results in comparison with older samples and Miocene volcanic episodes:

1. Zircon grains of 4–10 Ma from the active Cascade volcanic arc are rare in the Miocene rocks of the Tarheel and Empire formations. This suggests that during the Miocene, the Coos Bay depositional basin was not in connection with rivers draining the Cascades.

2. There is an increase in the abundance of 50–45 Ma zircon in the Tarheel and Empire formations. The simultaneous disappearance of arc-derived zircon and the reappearance of the 50–45 Ma retroarc Clarno-Challis zircon are best explained by intra-basinal recycling of the later from the underlying strata, likely from uplifted Tyee and Umpqua formation rocks along the Coast Range axis that were eroded and transported by the incipient Coos and Coquille rivers into the Coos Bay area.

Armentrout et al. (2020) interpreted the paucity of Cascade arc grains (<40 Ma) in the Tarheel and Empire formations to reflect late Oligocene to early Miocene uplift and exhumation of the southern Oregon Coast Range, effectively isolating the Coos Bay area (western limb of the “Tyee–Coos Bay”’ Cascadia Basin) from the Cascade volcanic arc source. This preliminary hypothesis is being further tested with modern river sand samples from the Coos, Coquille, Rogue, Umpqua, and Willamette Rivers to assess what zircon populations occur in each river basin.

Ilya Bindeman, University of Oregon Earth Sciences Department, and students are studying Pacific Northwest tuffs and lavas mapping the pathway of the Yellowstone plume and the timing of Cascade arc volcanism (Seligman et al., 2014). Mantle plume and volcanic arc magmas have different geochemical signatures. Five tuff samples from the Coos Bay area are being studied to determine their magma type and possibly the earliest onset of arc magmatism. The tuff layers occur in the Lower and Middle Coaledo Formation, Bastendorff Shale, Tunnel Point Sandstone, and Empire Formation (see Fig. 16). Associated zircon ages indicate that four of the tuff samples span 40.9–8.2 Ma., coeval with Cascade volcanic arc eruptions (Retallack et al., 2004). The fifth sample may date ca. 45 Ma but requires further stratigraphic study and geochemical analysis.

Seligman et al. (2014) and Camp and Wells (2021) propose pathways of the Yellowstone hotspot passing beneath the Coos Bay area, Siletzia, the Cascade Arc and across eastern Oregon and Idaho to Yellowstone National Park (Fig. 9). Camp and Wells (2021) show the hotspot track passing beneath the Coos Bay Basin between 40.5 and 40 Ma, during deposition of the Lower and Middle Coaledo Formation (45.0–39.8 Ma; Fig. 16).

I.N. Bindeman (personal commun., June 2021) has studied the five Coos Bay Basin tuff samples petrographically and under the SEM, and obtained bulk analyses using XRF for chemical analysis (at Pomona) and ICPMS for trace element analysis (at Washington State). The analyses were compared with Cascade rhyolite-rhyodacite and Yellowstone rhyolite and plotted on discrimination diagrams (Pearce, 1984) targeting trace elemental concentrations and ratios for arc-related vs plume-related distinct geochemical signatures.

Four analyzed Coos Bay Basin tuffs aged 40 Ma to 8 Ma have Cascade arc signatures suggesting that Cascadia arc volcanism goes back to 40 Ma. This 40 Ma volcanism matches the onset age of southern Oregon Cascade arc volcanism at 40.1 ± 0.6 Ma from the Fox Hollow tuff of the Fisher Formation near Eugene, Oregon (Retallack et al., 2004).

The fifth Coos Bay Basin sample from the lowermost Lower Coaledo Formation, sample 20COA17 (≤45 Ma), has a transitional geochemistry and is currently being further studied. This stratigraphically oldest sample could be of Yellowstone plume origin or suggest the earliest Oregon Cascade silicic magmatism.

Studies of detrital remanent magnetism (DRM) of the Cape Arago and Coos Bay area strata were published by Prothero et al. (2001a), Prothero and Donohoo (2001a), and Prothero et al. (2001b) that showed progressively more rotation back in time as have studies by Simpson and Cox (1977), Heller and Ryberg (1983), Wells (1989), and Brocher et al. (2017) for other Cascadia forearc rocks. These earlier studies provided stimulus for a restudy using current technology and rigorous testing of the DRM signature related to rock type. The current Coaledo Project study, led by Ray Weldon, University of Oregon Earth Sciences Department, and Scott Bogue, Occidental College, also expands the database down section in the lowermost Coaledo member and into the underlying Beds of Sacchi Beach. To date, 161 drilled sites have yielded 469 cores. More than 708 one-inch samples have been analyzed. Preserved remanence is highly variable but typically comprises a large, low-coercivity overprint that masks a smaller component more resistant to high alternating field (AF) and stepwise thermal demagnetization greater than ≈ 350 °C. Preliminary Coaledo Project results, for samples restored to paleohorizontal, show a high-coercivity component in the Coaledo Formation pointing in two, nearly antipodal directions that are rotated clockwise 70° with respect to the expected Eocene field direction (R.J. Weldon and S. Bogue, written comm., May 2021 and July 2021). This is less than the earlier Prothero and Donohoo (2001a) estimate of ~100° estimates for the Coaledo Formation rotation but overlaps within uncertainty the 67°14° for the Tyee Formation in the Roseburg area to the east of Coos Bay (Wells et al., 2000, 2014). Wells et al. (2017) include the Coos Bay and Roseburg areas in the same forearc “tremor” block which suggest they might have similar rotation histories.

Coaledo Project Team member Scott Bogue, Occidental College, has assembled a significant database on anisotropy of magnetic susceptibility (AMS) as a proxy for sandstone paleocurrents of the Coaledo Formation (see Gridler, 1961). Results from 62 stratigraphic levels with 230 cores have been studied (S. Bogue, written commun., June 2021 and July 2021). The anisotropy was determined by measuring the magnetic susceptibility of oriented core samples in nine orientations. Samples with internally consistent anisotropy were oriented to approximately depositional horizontality and statistically evaluated within a site and between sites. The results are widely dispersed as expected for wave deposited facies but provide a mean paleocurrent axis of 174°/354°, slightly counterclockwise of present day north-south. The mean AMS axis agrees well with paleocurrent measurements at several AMS sample sites, and with paleocurrents for Coaledo sedimentary rocks in other studies, all averaging a slightly west of north (present-day orientation) (Dott, 1966; Bird, 1967; Rooth, 1974; Ryberg, 79; Chan and Dott, 1986).

Combining the DRM and AMS data discussed here suggests that paleotransport of the middle Eocene Coaledo Formation was essentially north-northwest in agreement with earlier studies and that counterclockwise rotation of ~70° would have the middle Eocene Cape Arago Coaledo Formation deltaic system flowing approximately west-northwest 40 million years ago.

Our visits to the Cape Arago outcrops will provide us with observations of the strata and structural consequences of this multifaceted and episodic deformation of the Cascadia forearc.

NOTE: Latitude and longitude coordinates are from Google 2021, and may differ slightly from specific road locations, field stops and outcrop sites due to interference from cliffs and trees.

Drive to Cape Arago State Park at the south end of the Cape Arago Highway traveling from either North Bend or Coos Bay. The route to the state parks is well marked. Figure 20 outlines the field stops along the Cape Arago highway.

Park in the available spaces on the western edge of the loop roadway and walk southwest to the stone observation area at the western edge of the Cape overlooking Middle Cove on the south.

This field guide provides an overview of the sedimentary rocks of the Coos Bay area. Figure 16 provides the stratigraphic column with the age of formations based on recently obtained detrital zircon dates interpreted by Michael Darin. Figure 20 shows the geologic map of the field-trip area with the location of outcrop sites.

The Cenozoic rocks crop out on the west and east limbs of the South Slough Syncline, a fold developed during a late Oligocene to Miocene deformation (Fig. 19). The Beds of Sacchi Beach, Coaledo, Bastendorff, and Tunnel Point formations were deposited within the Cape Arago part of the Coos Bay area with conformity. Discontinuities within this conformable coastal succession are channel-scale. Regional unconformities that are coeval with the Cape Arago Paleogene occur farther east toward the folded and faulted area of the Siletzia suture zone southeast of Roseburg (Fig. 10) (Baldwin, 1973; Baldwin et al., 1973; Ryu et al., 1992, 1996; Wells et al., 2000).

Here at Cape Arago, the middle Eocene sedimentary rocks of the Beds of Sacchi Beach and Coaledo Formation are folded into a north-trending anticline, which is now truncated by a coastal terrace (Figs. 19 and 20). The anticline extends from South Cove to North Cove and is cut by a normal fault with down dropped strata on the west. A secondary fault trends through Middle Cove and may intersect the primary fault in North Cove. Resistant sandstone of the downfaulted Coaledo Formation form the seaward face of Cape Arago with deeply eroded fractures trending offshore.

North Cove and South Cove are eroded into the Middle Eocene Beds of Sacchi Beach (Wiley et al., 2015). At Cape Arago, the Beds of Sacchi Beach are gradationally overlain by the Coaledo Formation, but inland the Coaledo unconformably overlies early Eocene Tyee formation and Umpqua Group strata (Dott and Bird, 1979; Wiley et al., 2015). Sandstones of the Coaledo Formation contain mollusk fossils, particularly at Middle Cove Cape Arago, where sand dollars are moderately abundant, and at Sunset Bay (Turner, 1938). Mudstone interbeds of both the Beds of Sacchi Beach and the upper Lower Coaledo Formation contain foraminifera and ostracods (Bird, 1967; Rooth, 1974; Dott and Bird, 1979).

Walk eastward to the parking area by the Cape Arago restroom and then directly south across the road to the cliff top overlooking South Cove.

South from Cape Arago, the coastline follows the cliffs of the Seven Devils southward along Sacchi Beach (1st Point), Agate Beach (2nd point), and Merchant Beach to Fivemile Point (3rd point). Beyond Fivemile Point is the long sandy area near Whisky Run and Bullard’s Beach and finally the rocky headland of Coquille Point at Bandon. On a very clear day, Cape Blanco can also be seen far to the south. The eroded surface above the Seven Devils cliff is the Whisky Run Terrace, dated at 80,000 years (Figs. 21 and 22 and marine terraces data) (Griggs, 1945; McInelly and Kelsey, 1990).

Fivemile Point consists of vertically to overturned turbidite beds of the Sandstone of Fivemile Point, immediately south of the Fulmar fault zone. This fault zone is considered a southwestern boundary of the accreted Siletzia terrane (Wiley et al., 2015). Where we stand at Cape Arago, we are within the accreted Siletz terrane that extends north to southern Vancouver Island, British Columbia.

The Seven Devils cliffs along the eastern bluffs of South Cove expose a series of marine channel complexes grading upward from predominantly silty Beds of Sacchi Beach to increasingly sandy facies and conglomeratic facies interpreted by Dott (1966) and Dott and Bird (1979) as basal Coaledo Formation strata (Fig. 25). The Beds of Sacchi Beach are interpreted as slope facies deposited seaward of a delta in water depths of 100–200 m (Bird, 1967; Dott and Bird, 1979). The overlying lowermost Coaledo strata are interpreted as distal delta front facies (Dott, 1966; Chan and Dott, 1986).

Stops 3–8 follow an upward progression of channel facies similar to but sandier than those exposed in the Seven Devils cliff above South Cove (Fig. 25). This upward progression of shallowing facies suggests basinward progradation of sediment with progressively shallower water environments of deposition. This facies progression is how prograding clinoforms imaged by geophysical reflection techniques are interpreted (Fig. 26). The geometry of clinoforms represent three geometric facies: relatively horizontal topsets interpreted as shallow environments; foresets more basinward where sediments are deposited in progressively deeper water and most likely are finer grained; and farther basinward to bottomsets where downslope gravity-flows transport predominantly very fine-grained sediment and subordinate coarser-grained sandstones as submarine fans. The muddy-to-sandy channel facies of Seven Devils cliff are inferred to be upper slope foreset facies shallowing to outer shelf topset facies, as shown on Figure 26 for the “Approx. position of Fig. 25” label. Stops 3–10 are a progression of facies deposited in progressively shallower environments as the Eocene shoreline and shelf margin prograded into the deeper waters of the Coos Bay area. We will be traversing the environments along the transect of the dashed arrows of Figure 26.

Each of the depositional cycles we are to visit represents a depositional cycle called a parasequence in the nomenclature of sequence stratigraphy. We will explain this further at Stop 8. Each of the layers of seismic topset-foreset-bottomset in Figure 26 are a proxy for one parasequence. Figure 27 illustrates the parasequence framework for the Lower Coaledo, depositional cycle PS-1 to PS-14, and the progression of depositional settings interpreted from the rocks. We will relate each field stop to both this Figure 27 parasequence framework and to the clinoform diagram and seismic profile of Figure 26.

Depart the South Cove overlook, walking back past the parking areas to the trail to North Cove by the large information sign. Along the trail to the viewpoint, turn right and descend the trail taking the right turn down to North Cove. Anticipate that this is a marginally maintained trail and the bottom is eroded by waves resulting in a 2 m (6+ ft) muddy scramble down to the cobble beach. Traverse east (right) and then northward along the east cliffs noting the sandstone beds with abundant mudstone rip-ups.

The cliff area starting just before the waterfall and continuing north past the Simpson Reef Overlook exposes a series of submarine channel complexes interpreted to be an upper slope prodelta setting (similar depositional setting to upper forests of Fig. 26). The channel complex exposed at the waterfall (Fig. 28C) is ~100 m wide and 25 m deep (Dott and Bird, 1979). Careful correlation of layers along the North Cove cliffs and tidal platform exposures, using truncated stratal patterns, results in interpreting more than nine channel complexes, each composed of smaller-scale multistoried (stacked) channel and sheet facies. These Beds of Sacchi Beach channel complexes are eroded into the laminated prodelta slope facies exposing sedimentary fine-scale structures including flame, load-cast, debris-flow, symmetrical ripples, and traction ripple marks. The multistoried strata in-filling the channel complexes range from predominantly mudstone with laterally discontinuous slightly graded sandstone bearing abundant mudstone clasts, to trough cross-bedded sandstones, mostly fine-grained. As described by Dott and Bird (1979) the mudstone-filled channels typically have symmetrical bottoms (Figs. 28B and 28C) whereas channels filled with sandstone have more irregular margins (Fig. 28A–Coaledo Facies).

The basal channel complex scour formed by high-energy distributary flow (probably floods) that eroded upper slope channels as the high-energy and probably coarser-grained sediment-flow bypassed into deeper water. As the flow energy waned, the eroded channel began to fill with sediment. Subsequent erosion-bypass-fill cycles resulted in the multistoried channel-complex fill of at least nine such cycles being formed in the North Cove stratal section. John Armentrout and Dave Blackwell of the Coaledo Project Team are working with Christine Rossen, ExxonMobil retiree, who has worked on slope channel systems using seismic and well data. Rossen’s contribution is a detailed understanding to the geometry of slope channel systems using 3-dimensional seismic data. Our interpretations of the channel system using Google, drone, and outcrop photos provides clear definition of the succession of at least nine channel-form depositional geometries, some below and others above an erosional surface (“Tilted-Discordance,” Fig. 29) (C. Rossen, written commun., July 2021).

In the north-facing cliff slightly east of the cliff-top Simpson Reef Viewpoint, the tilted strata of the tidal flat can be observed to be truncated ~4 m (13 ft) up into the cliff (Fig. 29). This truncation surface correlates below 6, possibly 7 of the nine-channel complexes recognized in North Cove. The Coaledo Project Team working in North Cove suggests this truncation surface represents a slope-failure event, a submarine slope-slide, explained on Figures 29 and 30. This is a working hypothesis undergoing further study.

Before beginning the traverse out of North Cove, try to view the far cliff face of the cove just north of the Simpson Reef Viewpoint (fenced area at top of cliff). The next point, Hidden Cove Point, exposes the base of the Coaledo Formation, a yellowish concretionary sandstone immediately above the tidal platform with an overlying grayish laminated siltstone channel complex (Beds of Sacchi Beach–like facies). We will view and discuss this outcrop from Stop 4 at the Simpson Reef Overlook.

Return to the Cape Arago parking lot and drive north to the Simpson Reef Overlook parking area.

After enjoying the view of Shell Island, listening to the barking sea lions, and reading the kiosk information, move to the right along the railing to view into Sea Lion Cove looking northward toward the far cliff below Hidden Cove Point (Fig. 28A).

The cliffs along the north and east of Sea Lion Cove expose a base-of-cliff yellowish-orange conglomeratic and concretionary sandstone overlain by grayish laminated siltstone with curvilinear bedding characteristic of a channel complex. This stratal succession is the basal Lower Coaledo facies of the Coaledo Formation defined as conglomeratic sandstone containing coalified wood (Fig. 28A), and in some outcrops, thin coals or carbonaceous woody plant fragments. Coal beds have not been observed in either North Cove or South Cove at Cape Arago. Detrital zircons extracted from the basal Coaledo sandstone provide an approximate age of 45.0 Ma (Fig. 16).

Overlying this 7 m (23 ft) conglomeratic sandstone are 20 m (65 ft) of thin-bedded silty mudstone similar to what is mapped as Beds of Sacchi Beach. Beyond the corner of Hidden Cove Point, within Hidden Cove, this siltstone pinches out and fine-to-medium and coarse-grained sandstones persist along the coast beyond the farthest observable rocks observed from this viewpoint. The basal Coaledo sandstone is the precursor to the subsequently sandstone-dominated Lower Coaledo Formation facies all the way to Simpson Beach Cove at Shore Acres State Park. Based on this demarcation of a Lower Coaledo Formation datum, the Lower Member approaches 570 m (1870 ft) in thickness and the composite Coaledo Formation approximates 1755 m (5770 ft) in thickness. In fact, each of the depositional intervals (members, cycles, parasequences) vary in thickness from section to section along the outcrop exposures due to lateral facies variations.

Return to your cars and drive north ~0.64 km (0.4 mi) on Cape Arago Highway to the parking area near to the trail head for the Cape Arago Pack Trail (leads to a World War II observation site ruin) (43.320209°N, 124.385582°W). Across the road (NW) is a trail with barrier posts that leads to Stops 5 and 6. Hike west on this main trail. Just after exiting the spruce forest watch on the right for a heavily eroded trail heading toward the sea cliff. This trail takes you to Stop 5 at Ocean View Lookout.

You are standing on the Whisky Run Terrace above outcrops of the Lower Member of the Coaledo Formation. The view from this point provides one of the best sites to photograph storm waves crashing against the Shore Acres cliffs. Scramble down onto the rocks below the terrace. Figure 31 is a drone photo of this outcrop. Three of the four lowermost Lower Coaledo parasequences are exposed in this cliff face (PS-2, PS-3, and PS-4 of Fig. 27). The basal Coaledo parasequence, PS-1, is exposed farther south in Sea Lion Overlook Cove (Stop 4, Fig. 28A). The base of each parasequence is an erosional surface with deeply incised scours, some up to three meters deep. With very careful effort, it is possible to explore this outcrop downward into the uppermost part of PS-2 at the north end of the outcrop where fisher folk often cast their lines (Figs. 31A and 31B).

The depositional succession of PS-1, PS-2, PS-3, and PS-4 in the Lower Coaledo consists of fining-upward facies of amalgamated channels overlain by tabular beds. The channels are meter-scale, trough cross-bedded, medium-grained, well-sorted sandstones with abundant mudstone ripups (Fig. 31A). The tabular beds at Stop 5 are bedsets of 2–3-cm-thick sharp-based beds of fine-grained sandstone grading to silty mudstone, interpreted as turbidity-flow deposits. Bedsets of these turbidites are separated by decimeter-thick silty mudstone (Fig. 31B). Interspersed within this succession are thin asymmetric ripple sandstones, centimeter-scale contorted beds, load-casts, flame structures and small meter-scale slumps. Sandstone-filled burrows are present but rare and most dark laminae are covered with serrated plant debris, mostly finely broken leaf material. No molluscan fossils have been recovered from these facies. This facies succession is interpreted as distal delta-distributary or uppermost slope (foreset) submarine channel facies.

The base of PS-4 is easily examined along the upper part of the Ocean View Lookout section starting toward the north end of the outcrop (Fig. 31C: follow the dashed pale-green datum). Locate the truncated surface underlying the cross-bedded sandstone (Fig. 31C). This location is just across (north) of a smooth eroded surface bored with “cup-shaped” erosional remnants of rock-boring pholad mollusks (Evans, 1967). Find the base of the cross bedding and follow the surface south along the outcrop. Southward, the surface rises from the deepest incision to higher levels. Along this surface is a mottled pattern of Teredolites ichnofabric resulting from small mollusks that burrowed a log about two meters long (Bromley et al., 1984). Modern Teredo navalis are small elongate “worm-like” clams that bore into wood, making elongate burrows, and are known as “ship boring worms” or “termites of the sea.”

Walk along the sandstone surface tracing the base of PS-4 to the far southwest end of the outcrop and view the cliff face in the cove south of Ocean View Lookout (Fig. 31). Most of the back wall cliff is PS-4, a parasequence of massive, amalgamated sandstone channels. The erosional base of this thick amalgamated sandstone is the continuation of the basal erosional surface of PS-4 where the teredo-bored log was observed.

The succession of basal erosion, amalgamated cross-bedded channels and tabular bedding shows the character of the four lowermost Lower Coaledo parasequences, PS-1, PS-2, PS-3, and PS-4 (Fig. 27). The drone photo of Figure 31 shows the deep incision of PS-3 into PS-2 and significant lateral thickness changes that make lateral correlation of parasequences challenging even with the continuous drone photographic survey.

Return to the top of the outcrop and follow the well-worn cliff-top trail northward toward Collapse Cave Point, the southwest boundary of Simpson Beach Cove. While traversing this area, observe the succession of depositional facies exposed in the cliffs, but beware of cliff-edge overhangs of the unstable soils upon the Whisky Run Terrace. When you reach a point that the trail seems to end back up a few meters, take the trail through the Salal bushes for several meters to discover an open wooded area, and proceed north-northwest to the point forming the southwest margin of Simpson Beach Cove.

The informal name for this point comes from a large sea cave that collapsed in the 1980s leaving a large funnel-shaped debris field of chaotic boulders. Looking across Simpson Beach Cove provides a view of the south facing cliff along which Margie Chan defined the lowermost cycles 6–10 in her 1982 dissertation, subsequently published in-part as Chan and Dott (1986).

From the top of Whisky Run Terrace, carefully descend the west-facing sandstone benches to observe the erosional base of PS-5 (see Fig. 32 for this observation station red-square A).

PS-5 is dominated by amalgamated cross-bedded coarse-grained sandstone channel and an upper facies of medium-grained trough cross-bedded channel deposits separated by decimeter thick silty mudstone through which numerous sandstone dikes were extruded. There are essentially no centimeter-scale turbidites observed at Stop 5. Higher in the section just left (north) of the dramatically truncated sandstone beneath an overhanging sandstone (see Figure 32C for location) occurs the first fair-weather oscillation wave-ripples suggesting shallowing compared to the Ocean View Lookout exposures of PS-2 to PS-4 (R.J. Dorsey, written commun., May 2021).

Three other features worth comment at this outcrop are: (1) Ophiomorpha burrows, 2–3 cm dark rings of pellets used by shrimp-like organisms to reinforce their burrows; (2) fossilized wood either coalified or petrified reflecting very different preservation processes; and (3) the dense fracture pattern within concretions compared to the enclosing sandstones due to the increased brittleness of the carbonate and iron cementation (“Eschelby” Joints).

PS-5 illustrates a lateral facies change from the Ocean View Lookout succession of thick amalgamated channels and tabular bedding with lateral facies changes that we will observe more strongly developed in the more northerly and stratigraphically higher outcrops. The drone photo (Fig. 32—dashed light blue rectangle) captures this lateral change from lower thicker coarse-grained channel facies upward and laterally to meter-scale tabular trough cross-bedded sandstones separated by mudstone.

Parasequence 5 in Figure 32 is the best example of this lateral facies gradation. At the far right just above the yellow cycle boundary are amalgamated channel facies that grade left and upward into the tabular bedding facies.

Farther north across Simpson Beach Cove and toward Bathers Cove the basal amalgamated coarse-grained sandstones are laterally replaced by mudstone. These lower parasequence silty-mudstones of PS-6 to PS-14 cycles thicken with upward increasing numbers of interbedded hummocky cross-bedded sandstone within storm-dominated shoreface successions (Figs. 33 and 34). Our preliminary interpretation of the above lateral and vertical facies pattern suggests longshore-wave transport of sediment northward from distributary channel mouth bars along the shoreline to shoreface depositional settings. This pattern is envisioned as similar to the distributary mouth bar at the mouth of South Pass of the Mississippi River delta system. This depositional setting is a preliminary interpretation needing additional analysis. Our team does not have consensus on the depositional setting or paleowater depths.

Each of PS-6 to PS-14 farther north along depositional strike are coarsening-upward and shoaling upward shoreface facies. As a working hypothesis, Armentrout infers a “fixed in-place” distributary system due to subtropical vegetation “rooting” the distributary margins and a strong wave influence transporting most of the delta front sediments northward by strong longshore drift (present day orientation) (Li et al., 2011). This facies distribution suggests an asymmetric delta complex with a strong fluvial distributary system for the strata from North Cove to Sunset Bay. Future research must include Lower Coaledo distributary systems along Seven Devils cliffs and at Agate Beach headland in the paleogeographic analysis. Because of the marine terrace truncation of the section and complete absence of significant deposition dip exposure this inferred deposition analysis is very tentative. Additionally, the channel complexes exposed along the Seven Devils cliff area must be included in the analysis (see Fig. 25).

Back-track to Ocean View Lookout, returning to the left and retrace along the main trail to the entry point on Cape Arago highway. The next stop is in Simpson Beach Cove, accessed either on a northward trail slightly west of the highway on the trail to Ocean View, or you can return to your vehicle and drive into Shore Acres State Park and take the south trail to Simpson Beach Cove.

Six parasequences, PS-3 to PS-8, are exposed along the north side of Simpson Beach Cove; PS-5 to PS-8 are accessible from the beach. Figure 33 shows the section and the correlation to the Chan and Dott (1986) measured section.

The Simpson Beach Cove section exposes the progressive shallowing of shoreface parasequences first observed at Collapse Cave Point in PS-5 (Stop 6), where the occurrence of fair-weather oscillation wave ripples (Fig. 32C) suggests shallower offshore channel complex facies than the deeper water channel complexes of Stops 3, 4, and 5.

The parasequences in Simpson Beach Cove from PS-4 to PS-8 coarsen upward with increasingly thick intervals of massive and laminated mudstone and increasing thicknesses of hummocky cross bedding. Each is capped by upper shoreface trough cross-bedded and parallel-laminated medium-grained sandstone. This facies succession fits the shoreface depositional model of Pemberton et al. (2012). This same facies succession is appropriate for PS-10 to PS-14 observed at Sunset Bay. Hummocky cross bedding is interpreted as storm-wave event deposition (Dott and Bourgeois, 1982).

Parasequence PS-7 mudstones yield the first significant molluscan fossils of the Protobranch Community (Hickman, 1984). This molluscan community of infaunal deposit feeders such as Nuculana, infaunal suspension-feeding Pitar, and carnivorous gastropods are characteristic of organic-rich offshore mudstone of middle neritic to upper bathyal depths. Nesbitt (1995) included this community in her Nuculana assemblage recovered from prodelta facies within the middle Eocene Cowlitz Formation in southwestern Washington. This fauna occurs in the silty mudstones of PS-7 to PS-14 shoreface facies, often in association with hummocky cross-stratification. The co-occurrence of the Protobranch Community mollusks with the fine-grained hummocky cross-stratification affirms the paleowater depths of lower to middle shoreface facies of middle to outer shelf/upper slope settings, probably in water depths less than 200 m.

Southward across the beach, the south Simpson Beach Cove section provides access to the uppermost PS-7 trough cross-bedded facies and the shoreface succession of PS-8. In the uppermost trough cross-bedded sandstone of PS-8 is a small distributary channel nested within the uppermost trough cross-bedded to parallel-laminated facies overlain by the fossiliferous laminated mudstones of PS-9.

There is a large beach boulder offset from the south cliff face near the small distributary channel that exposes a complete, bi-directional downlapping hummocky cross bed. Hummocky cross-bedded sandstones are interpreted as storm wave deposits (see Dott and Bourgeois, 1982, and references therein). Peter Ruggiero, Oregon State University Oceanography, has modeled the expected water depths of sediment movement creating hummocky sand beds for coarse silt to medium sand using 1980–2015 buoy data for wave energy (Fig. 34: written comm., Feb. 2020). Ruggiero’s model suggests storm wave base could be as deep as 270 m (886 ft) for exceptionally large storm waves but is more likely to be ~170 m (560 ft) of water depth. This suggests that the Lower Coaledo Formation shoreface parasequences PS-6 to PS-14 represent shelf sedimentation cycles driven by either delta evulsion events or eustatic cycles. There are similar “hummocks” in the uppermost channel complex facies of the Beds of Sacchi Beach within the Sea Lion Overlook Cove of North Cove, Cape Arago (below Stop 4). Dott and Bird (1979) observed these hummocks and attributed a paleowater depth based on biofacies analysis at 100–200 m (300–600 ft) (Dott and Bird, 1979). This interpretation would suggest the upper Beds of Sacchi Beach channel complexes occurred in a much shallower basin-margin slope setting rather than a deep-water continental margin slope setting, an interpretation compatible with the interpreted Paleogene setting of a shelf basin landward of the initial accretionary prism outer shelf high (see Fig. 12A and Fig. 18).

Depart Simpson Beach Cove and go to the main parking area of Simpson Cove State Park (43.324232°N, 124.387047°W). Restroom facilities are available along the parking area.

Our traverse to Stop 8 starts from Shore Acre State Park’s cliff-top gazebo. The exhibits posted inside the gazebo are worth some time learning about the history of the Simpson family estate. Leave the gazebo taking the cliffside trail north past the cement slabs of the old Simpson estate tennis court. At the north end of this area is a cliff top view into the large cove called Bathers Cove.

The stratigraphy along the north (far) margin of Bathers Cove exposes shoreface facies of PS-6 to PS-9 (Fig. 35). The western cliff is eroded in PS-6; PS-4, and PS-5 occur farther west and are accessible with difficulty only at very low tide.

The Bather’s Cove parasequences sandstones of PS-7, PS-8, and PS-9 are finer grained than the sandstones at Collapse Cave Point and Ocean View Lookout, perhaps due to either waning energy of longshore drift or decrease in grain-size within the distributary source area. Within Bathers Cove, exposures of PS-7 and PS-8 have thick intervals of silty mudstones and meter-thick hummocky cross-bedded fine-grained sandstones. Figure 35 provides a photo with facies identification for each parasequence. The siltstones of PS-7 and PS-8 are relatively massive to laminated and fossiliferous with the prodelta to offshore Protobranch Community mollusks discussed at Stop 7, Simpson Beach Cove. The massive bioturbated to parallel-laminated medium-grained sandstone capping PS-8 contains thick-shelled Venericardia shells, suggestive of upper shoreface environments.

Our next stop at the Lighthouse Viewpoint can be reached by hiking along the coastal trail or returning to your vehicle and driving north on the Cape Arago Highway to the parking area at the Lighthouse Viewpoint. This viewpoint is before (south of) Norton Cove and Sunset Bay.

The cliff top view from this Cape Arago Highway pull-out is across to Qochyax Island and the Cape Arago Lighthouse constructed on Chief Island that extends northwest as Gregory Point (Fig. 36). The northwest curvature of Gregory Point suggests perhaps the northern nose of the anticline associated with the Cape Arago fault (see map Fig. 20). Both Qochyax and Chief Islands, previously U.S. Coast Guard properties, are now private property of the Confederated Tribes of Coos, Lower Umpqua, and Siuslaw Indians.

Qochyax Island years ago was a beautifully forested island. Nesting sea birds altered the chemistry of the soils with their fish-rich droppings, resulting in the demise of nearly all the vegetation (R. Frey, personal commun., July 2018).

The strata of the cove below the Lighthouse Viewpoint are PS-11 and correlate with the strata between Qochyax Island and the mainland sea cliffs of Sunset Bay observed to the north. Figure 36 shows the correlation of parasequences from Bathers Cove to Lighthouse Viewpoint and across Sunset Bay to Qochyax Island. Parasequence numbers are placed at the top of each coarsening-upward parasequence. Structural dip is to the east. This correlation incorporates work by Allen and Baldwin (1944), Rooth (1974), Ryberg (1978), Chan and Dott (1986), and Coaledo Project Team’s correlations using drone photos. The deformation pattern of faults and fractures observed in each Coos Bay area time-stratigraphic unit is summarized in Figure 37. The Coos Bay Basin Cape Arago section deformation phases correlate with the deformation phases for the Paleogene of Siletzia as discussed in the tectonic overview in Section I of this document.

PS-11 differs from the other Lower Coaledo shoreface parasequences in having a thick, progressively deepening facies succession rather than a sharp transgressive flooding surface at the base (see Fig. 27). Additionally, the uppermost facies of PS-11 consist of conglomerate and coarse-grained sandstone with shallow-water mollusks and leaf fossils interpreted by Chan and Dott (1986) as fluvial. This is the shallowest facies interpreted in the Lower Coaledo and is overlain by progressively thicker mudstone-rich parasequences suggesting the onset of relative sea level rise proceeding the regional transgression of the Middle Coaledo member (Fig. 27). This progressive change in parasequence facies is the focus of Stop 10 at Sunset Bay.

Our next stop is Sunset Bay. Either continue north on the coastal trail or return to your car driving to and park along the shore of Sunset Bay. This is a beautiful setting for a multitude of perspectives including tectonics, ancient forests, intertidal platform erosion, paleontology, and lithostratigraphy.

Sunset Bay is an arcuate bay formed along a complex set of northwest-trending faults transverse to the strike of bedding (Fig. 36). The faults have a right lateral offset of ~140 m (460 ft) between the north and south side of the bay. Whether this displacement is along a single fault or a series of smaller faults such as those exposed along the bay margins is unclear. Orientation of drag folds suggests that the fault motion was right/lateral oblique slip (Ehlen, 1964). Neotectonic and geohazard studies in the Coos Bay area focus on Pleistocene marine terrace deformation (McInelly and Kelsey, 1990; Madin et al., 1995; Snavely and Wells, 1996; Wiley et al., 2015). The 80 k.y. B.P. Whisky Run terrace is elevated ~20 m above the modern intertidal terrace, having been uplifted at a rate of 0.55–0.78 m/k.y. (McInelly and Kelsey, 1990). The tilted beds of the Coaledo Formation are related to a late Oligocene–early Miocene phase of north-northeast–directed compression that uplifted and truncated the Eocene-Oligocene formations of the Coos Bay Basin most probably about <30 to >25 Ma (Fig. 37, Phase 1 deformation).

Along Big Creek on the south side of Sunset Bay beach, are broad, up to 8–10 m (26–32 ft) in diameter tree stumps rooted in the modern tidal platform. These root systems of spruce and hemlock are not “Ghost Forest” stumps drowned by land-subsidence from earthquake events as occur in many estuaries along the Oregon and Washington coast (Hart and Peterson, 1997). The more shoreward root systems can be seen to be rooted into a soil horizon with a peat layer developed on the truncated Coaledo Formation sediments. This prompted Armentrout (1980) to interpret that the trees had been growing on the Holocene flood plain of Big Creek and killed by sea water as the sea carved out the Sunset Bay amphitheater long after sea level attained its present position. This interpretation is strengthened by Hart and Peterson (1997) based on a much stronger database of 275 stumps at 14 localities from Coos Bay north to Neskowin, Oregon. That study found three mechanisms for regression and transgression of the surf zone and “drowning” of the trees: (1) coastal regression and removal of sand barriers; (2) eustatic change of sea level; and (3) vertical displacement of the Cascadia margin. Carbon 14 data on one root from along Big Creek yields an estimated age of ~1200 yr B.P. (Armentrout, 1980; Southern Methodist University Radiocarbon Laboratory Sample 593-B8/12-Count 1439).

Walk to the north end of the parking areas at Sunset Bay and enter toward the cliff area at the boat ramp (Fig. 38). This area is accessible across the intertidal platform only at low to minus tide conditions.

The intertidal platform of north Sunset Bay displays an intricate fault pattern frequently mapped by university students (Fig. 38). It is also the site of detailed studies of the origin of such rock platforms. Intertidal platforms are called wave cut surfaces. Retallack and Roering (2012) made detailed studies differentiating weathering process and sediment removal contrasting the cliff processes with those of the intertidal platform. They found that despite continuing tectonic uplift, the rock platform remained in the intertidal zone coincident with the water table. The shape and elevation of rock platforms in coastal and fluvial settings appear to reflect differences in strength between bedrock and saprock (weathered fractured bedrock) within the zone of water table fluctuation. They found that waves and floods do not “cut” the rock but remove clasts already weakened by weathering thus exposing the local water table as a bedrock surface. Such marine surfaces might more appropriately be called intertidal or water-table bedrock platforms. The intertidal platform is also weakened by rock-boring pelecypods Penitella and Adula. Their unoccupied borings in turn become a habitat for a succession of other organisms (Armentrout, 1975).

Parasequences PS-12, PS-13, and PS-14 are accessible along the north side of Sunset Bay at low tide (Fig. 39). The top of PS-12 and PS-14 form prominent points along the shoreline. PS-13 has a prominent point in mid-parasequence hummocky cross-bedded facies but the parasequence top is stratigraphically higher and somewhat obscured by cliff talus and vegetation. The top of PS-13 is a meter-thick parallel-bedded partially bioturbated fossiliferous sandstone exposed at the base of the cliff under overhanging trees. This PS-13 sandstone-top does not make a prominent point. The faulting between PS-13/PS-14 is along several splays making identification of unit thicknesses a challenge. A zircon sample from the uppermost PS-14 sandstone yields a preliminary age of 42.5 Ma (Fig. 16).

These three parasequences have progressively thicker and more extensively bioturbated mudstone intervals, suggesting a longer period of deep-water sedimentation or alternatively a more rapid rate of sediment accumulation as the sea floor subsided. Each parasequence is a shoaling-upward succession of mudstone with rare thin gravity-flow sandstone beds overlain by hummocky cross-bedded sandstone interbedded with mudstones and capped by trough cross-bedded and parallel-laminated fossiliferous medium to coarse-grained sandstone. These three muddier and more highly bioturbated parasequences are overlain by the deeper-water mud dominated Middle Coaledo that represents a basin-wide transgression. Our sequence stratigraphic analysis suggests the relative rise in sea level began with PS-12 as illustrated in Figure 27.

Sedimentary structures are well-developed and beautifully exposed in outcrops along the north side of Sunset Bay and on the cliffs and terraces beyond the southwest edge of the bay entrance (Dott, 1966; Ryberg, 1978; Dott and Bourgeois, 1982; Chan and Dott, 1986). Primary sedimentary structures include tabular, trough, wedge and hummocky cross-stratification, ripple cross-stratification, and rare flute and groove sole marks. Structures resulting from gravity deformation of the sediments include contorted bedding and flare, and ball and pillow structures. Very small clastic dikes and isolated sandstone “load-balls” represent liquification structures. Bioturbation includes both vertical and horizontal burrows of the skolithus ichnofabric.

Our next stop is Lighthouse Beach and Yoakam Point (parking at 43.339898°N, 124.359355°W). Access is limited to the trail marked Yoakam Point State Park ~0.8 km (0.5 mi.) north-northeast from Sunset Bay along the Cape Arago highway (see map Fig. 20). The entrance trail arrives at a deep gully with a rustic trail into a cove between two sandstone points of Upper Coaledo Formation sandstone. An excellent view of Lighthouse Beach is available by crossing the cliff top gully and walking on the Whisky Run Terrace to the northern point overlooking the access cove. At the far western end of Lighthouse Beach is Gregory Point formed on the upper sandstones of the Lower Coaledo member. Closer to the viewpoint are three prominent sandstone points within the Middle Coaledo. The viewpoint where you are standing is underlain by the lower most sandstone of the Upper Coaledo Formation (Fig. 40). Our facies analysis has only begun on the Upper Coaledo section and will lead to a parasequence analysis based in part on the excellent work of Chan and Dott (1986).

NOTE: Access to Lighthouse Beach and the Middle Coaledo section requires scrambling down the trail into the access cove and over the western point sandstone, which floods early with the rising tide. This sandstone point and two others of tabular turbidite sands along the Lighthouse Beach (Sites 3 and 6) are a concern for rising tide entrapment. All other access to Lighthouse Beach is through private property from the bluff above the beach, which requires special permission to use gated stairways.

The Middle Coaledo member is predominantly laminated mudstone with three intervals of pronounced sandstone. The Middle member base at the west end of Lighthouse Beach overlies a prominent cliff forming sandstone of uppermost Lower Coaledo (PS-14) (Fig. 40 Google image). Lowermost Middle Member strata consists of several hummocky cross-bedded sandstones of medium to coarse grain with broken mollusk shells interbedded in laminated silty mudstone. Several east-west faults offset the top of the Lower Coaledo at Gregory Point, making correlation of the Lighthouse Beach cliff face strata to the lower beach sea floor facies a challenge (Fig. 40).

The Middle Coaledo is ~760 m (2500 ft) thick and predominantly mudstones that have been correlated east and south in outcrop and wells and interpreted as a Coos Bay area-wide transgression. The silty mudstones are laminated and locally bioturbated. Foraminiferal faunas suggest outer neritic to upper bathyal paleowater depths (Rooth, 1974). Asymmetric ripples, small channels, debris flow facies and downslope detached bedding suggest a slope setting. Google Map 2021 images show sea floor areas of rotated blocks interpreted as slumps. Google Earth-Pro images of 5 May 5 2013, have even better low tide images of the modern shoreface and the rotated strata (Fig. 41).

Several volcanic tuff beds occur interbedded with the Middle Coaledo silty mudstone. Initial geochemistry suggests a volcanic arc magma source, most likely the Cascadia arc (I.N. Bindeman, written commun., June 2021). The Middle Coaledo tuffs are bracketed by zircon ages of 42–39 Ma (see Fig. 16). This correlates well with Wells et al. (2014) onset of early Cascade Fisher Formation volcanism at ca. 40 Ma (see Fig. 11) and Retallack et al. (2004). Stern and Dumitru (2019) report onset of Washington state Cascade arc volcanism as early as 45–42 Ma. Allen and Baldwin (1944) report tuff in the Lower Coaledo along the Coquille River. We have not observed any tuffs in the Lower Coaledo Formation Member within the Cape Arago section, but we have collected a tuff from lowermost Coaledo east of Coos Bay near the Coos River that is currently being studied (see Section II, Tuff Layer Analyses). If this tuff is found to be of Cascade Arc magma type, then onset of Oregon forearc volcanism would need to be reassessed.

Approximately 500 m (1600 ft) up Lighthouse beach is the first thick sandstone point consisting of tabular fine-grained sandstone beds, most with slightly scoured basal contacts and gradational tops (Fig. 40, Site 3). These deposits are interpreted as gravity-flow deposits, turbidites (Rooth, 1974). The dominantly fine- to slightly medium-grained sandstone precludes development of a full Bouma textural succession. Rare sole marks and flame structures indicate flow to the north northwest. These turbidite sandstones occur above an upward-thickening succession of silty to sandy beds, and below similar fining and thinning-upward beds, suggesting that a gravity-flow system was migrating across a sloping sea floor with a low-gradient that facilitated sand deposition.

A second prominent turbiditic sandstone point (Fig. 40, Site 6) occurs farther up section past a smaller point formed from a massive sandstone with abundant mudstone rip-ups that has been interpreted as a sandy debris flow deposit (Fig. 40, Site 5).

A large slump mass (Site 4) occurs between the lower turbidite (Site 3) and sandy debris flow (Site 5) points but is seldom exposed due to thick beach sands (Fig. 42). The 2021 Google Map image does “image” this slump deposit (Site 4) as a grayish “rectangle” just seaward of the shadow of the upper beach ridge.

The succession of mid to upper Middle Member rotated block, turbidite, debris flow and slump facies suggests a slope environment. We surmise that ca. 43–42 Ma (very preliminary) following the regional transgression of the lowermost Middle Coaledo, the sandy river system of the Coos Bay area was prograding rapidly toward the north (today’s orientation) with increased rates of sediment transported to the slope causing instability and slope failure. Only 150 m east from the upper turbidites, and 140 m stratigraphically higher, is the first of eight prograding parasequences of the overlying Upper Coaledo Member. The sixth Upper Coaledo parasequence includes a 2 m coal seam associated with estuarine fossils and conglomeratic sandstones. This facies association suggests the shoreline and coastal plain had reached the Eocene Yoakam Point area (Figs. 43–46). The Middle to Upper Coaledo contact is gradational from laminated siltstones to fine-grained hummocky cross-bedded sandstone and a parasequence capped with poorly developed fine- to medium-grained trough cross-bedded sandstones. This parasequence is the point we crossed when we entered Lighthouse Beach from Yoakam Point State Park (Fig. 43). A zircon sample from a meter-thick sandstone low in this parasequence yields a preliminary age of 41.5 Ma (Fig. 16).

The Upper Coaledo is ~425 m (1400 ft) thick and is exposed along the prominent sandstone points and deep coves eroded into mudstones between sandstone points (Fig. 43). There are seven significant and one subordinate parasequence cove-to-point successions between the base of the Upper Member and Mussel Reef, the most seaward prominent point. Cove 1 is the Yoakum Point State Park trail cove.

Upper Coaledo Formation facies are significantly different from the consistently coarsening-upward shoreface succession of the Lower Coaledo member. The stratigraphically lower Upper Coaledo parasequences have thin mudstone lower intervals and coarsen upward to fine- to medium-grained trough cross-bedded sands in the upper intervals. Higher in the Upper Coaledo the parasequences coarsen to coarse-grained and conglomeratic channel deposits. Most fossil occurrences are wave-sorted concentrations or dispersed broken shells. A two-meter coal bed has been correlated with the Beaver Hill coal seam of the Coos Bay Coal Field east of Cape Arago (Figs. 43, Cove #4, and Fig. 44). The coal is situated above a trough cross-bedded sandstone with Ophiomorpha burrows (Fig. 45). Overlying the coal is a 50 cm (1.6 ft) thick prograding fossiliferous sandstone bearing estuarine molluscan fossils (Rooth, 1974; Chan and Dott, 1986). This suggeststhat the coal-swamp environment was probably very close to the Upper Coaledo Formation late Eocene shoreline. Farther above the coaly facies, the section is dominantly conglomeratic channels and pebbly sandstone with shell fragments suggesting shoreline deposition adjacent to streams (Fig. 46).

Stratigraphically above Mussel Reef is a deep cove, locally known as Pirate’s Cove (Fig. 43). This cove is eroded into a thick mudstone interval overlain by the uppermost sandstone of the Upper Coaledo Member (Fig. 47). The mudstone interval may correlate with the peak of the middle Eocene Climate Optimum and we hope to sample some of the slump-buried mudstones for micropaleontologic confirmation of the inferred climate. In the meantime, we rely on a preliminary zircon date for the uppermost Upper Coaledo sandstone of 40.1 Ma for the tentative correlation of the uppermost Lower Coaledo with the middle Eocene Climate Optimum (Fig. 47).

NOTE: Mussel Reef, Pirate’s Cove, and the top of the Upper Coaledo can be reached walking south a half mile from the Bastendorff Beach Road, or by requesting access through the RV camp along the Miner Creek Valley.

Having departed Yoakam Point State Park and returned to your vehicle parked along Cape Arago Highway, drive northeast along the highway and turn left (north) onto Bastendorff Beach Road. Drive past the campground and down the hill, stopping at the base of the incline and parking on the road shoulder next to Bastendorff Beach (parking at 43.342883°N, 124.348324°W). The traverse to the only remaining Bastendorff Shale outcrops is southwest from the road along the base of the steep hillside looking upward until the outcrop is discovered nearly to the far end of the spruce forest. Access to the outcrop directly from the beach is obstructed by Miners Creek and wetlands.

The Bastendorff Shale (Madin et al., 1995) of late Eocene to earliest Oligocene age encompasses a 950 m (3100 ft) stratigraphic interval between the uppermost Coaledo Formation sandstone at Pirate’s Cove and the lowermost Tunnel Point Sandstone. The lowermost 200 m (600 ft) of probable Bastendorff Shale is covered by alluvium along Miners Creek valley. Weaver (1945), Allen and Baldwin (1944) and students from the University of California at Berkeley in the early 1950s measured, described, and collected fossils from the Bastendorff mudstone when the outcrops were relatively unvegetated sea cliffs (Tipton, 1975).

Construction of the Coos Bay entrance South Jetty in 1926 resulted in rapid growth of Bastendorff Beach (Fig. 48). Now the old sea cliffs are completely vegetated except for a single cliff-face outcrop within the lower Bastendorff Shale stratigraphically above the sandstone in Tipton’s measured section (1975, Fig. 6) (Fig. 48A). Detrital remnant magnetism (DRM), micropaleontological, and tuff samples have been collected from this site.

There is in fact a second Bastendorff “outcrop” ~30 m (90 ft) farther southwest along the bluff where a heavily vegetated “rib” extends across the forest floor, approximately at the boundary of the evergreen spruce forest and deciduous trees. This rib is underlain by the lower Bastendorff Shale sandstone noted in Tipton (1975). Pickax excavation uncovered the sandstone. Zircon and DRM samples were collected. The zircon age is 37.8 ± 0.7 Ma (Fig. 16).

There are three lithofacies described for the Bastendorff Shale (Tipton, 1975). Most of the formation is laminated to thin-bedded mudstone with some tuff beds. The tuff sampled in Figure 48A has a geochemistry correlated with Cascade arc magma type (I.N. Bindeman, written commun., June 2021). Fossils include a middle to lower bathyal foraminiferal assemblage and thin-shelled small deep-water pectens (Tipton, 1975). The several tabular sandstone beds in the lower part of the measured section contain a fauna suggestive of neritic depths. These sandstones, visited by Armentrout in 1965, were tabular sharp-based medium- to coarse-grained sandstones with silty mudstone interbeds. The sandstone contained several fossil leaves that were rolled-up like cigars. It is suggested that these lower Bastendorff sandstones are deep-water turbidites with associated displaced neritic foraminifera. The third facies of the Bastendorff are uppermost tuffaceous mudstone with rare molluscan fossils in increasingly silty and sandy intervals. The Bastendorff Shale is conformably overlain by the Tunnel Point Sandstone.

Drive northeast along Bastendorff Beach Road to the public restroom building and parking lot. Tunnel Point is the bluff you just passed and now is toward the south. The Tunnel Point Sandstone outcrops are along the forested bluffs between the basal sandstone point and Coos Head Road.

The Oligocene Tunnel Point Sandstone (Madin et al., 1995) occurs as a truncated wedge between Coos Head Road, Bastendorff Beach Road and the underlying Bastendorff Shale (Fig. 48). It is unconformably overlain by the late Miocene Empire Formation (Schenck, 1927; Weaver, 1945). This relationship is complicated by Coos Head fault mapped along “Gold Washer’s Gully” where Coos Head Road is constructed (Madin et al., 1995). Descriptions of the contact are reviewed by Weaver (1945) from field mapping and early published papers based on the section when it was still a sea cliff exposure. The angular unconformity above the Tunnel Point Sandstone may represent the onset of Coast Range uplift but this possibility is still under study (see Figs. 16 and 19, and Section II, Coos Bay Basin Tectonic Setting: Basin Isolation Hypothesis).

The basal sandstone of the 245-m- (800-ft-) thick Tunnel Point Sandstone forms the strike wall at the contact between the Bastendorff Shale and the Tunnel Point Sandstone (Fig. 48B). The contact is interpreted as conformable. Today that contact is buried by a yard-debris cone at the base of the cliff. The basal sandstone is fossiliferous and concretionary. Up section from this facies, the Tunnel Point Sandstone, consists of poorly to tabular-bedded fine-grained tuffaceous sandstone and sandy siltstone. The molluscan fauna of 33 species (Weaver, 1945) correlates with the Eugene Formation in the southern Willamette Valley (Hickman, 1960; Retallack et al., 2004).

Prothero and Donohoo (2001b) report clockwise rotation of the Tunnel Point Sandstone. They report a rotation of 106° ± 18°, consistent with their 105° ± 5° measured from the Coaledo Formation (Prothero and Donohoo, 2001a). New paleomagnetic data from the Cape Arago Coaledo Formation section suggests approximately 70° ± 25° of post Eocene clockwise rotation. Continuing analysis of Cape Arago strata, including the Tunnel Point Sandstone and Bastendorff Shale will reassess the rotation history using current methodologies.

Tunnel Point received its name from the occurrence of a large sea cave within the main point of the lower sandstone (Fig. 48C). The cave is 36 m (120 ft) deep, 15 m (50 ft) wide, and 7 m (23 ft) high. The sea cave was eroded along a fault and expanded in soft shaly sandstone with a northeast-facing entrance still accessible by “bush-whacking” along the cliff base. Blocks of sandstone frequently fall from the fractured ceiling and extreme caution is urged if entering the cave by way of a muddy crawl.

The next stop is at the Coos Bay entrance south jetty where it abuts the Empire Formation sandstone at Coos Head. Continue driving north on Bastendorff Beach Road and turn left onto Coos Head Road driving to south jetty (Fig. 48). If the waves are gentle, scramble onto the concrete core of the jetty and walk to the right toward the southeast end.

The jetty was constructed in the 1920s using blocks of volcanic breccia from quarries east and north of Coos Bay where local exposures of the Siletzia terrane basalts are mapped as Roseburg volcanics. More recently blocks of blue schist have been sourced from accreted terranes near Bandon. The basalts were transported to the mouth of Coos Bay on wooden tramways around Coos Head and through tunnels cut through the Empire Formation sandstone (Figs. 48D and 48E). The 1928 completion of the South Jetty resulted in the accumulation of sand between Mussel Reef at Yoakam Point and the South Jetty creating Bastendorff Beach.

Across the Coos Bay entrance and beyond the North Jetty is the beginning of the Oregon Dunes National Recreation Area. The dune field extends nearly 80 km (50 mi.) north to the city of Florence, an area of 16,187 ha (40,000 acres) and is the largest expanse of temperate coastal dunes field in the world (Cloyd, 2018). The Coos Bay–Cape Arago region at the south and the Heceta Head and Cape Perpetua area north of Florence are uplifted areas due to differential strain resulting from compression within the Cascadia forearc. The topographically low area between the uplifts facilitates accumulation of sand.

The late Miocene Empire Formation consists of massive bioturbated fossiliferous sandstone with intervals of abundant fossiliferous concretions. The Coos Head section on the southwest limb of the South Slough Syncline exposes 410 m (1310 ft) of Empire sandstone.

This section of the Empire Formation unconformably overlies the Oligocene Tunnel Point Sandstone (Weaver, 1945) (Figs. 16 and 20). Missing from the Coos Head–Tunnel Point area section is the Tarheel formation (“Miocene beds” of Madin et al., 1995) that unconformably underlies the Empire at Stop 19. Both the Tarheel and Empire formation strata crop out along the South Slough Syncline formed during late Oligocene–early Miocene SW-NE compression that uplifted the Paleogene section. The SW-NE compression was related to a change in the vector of Cascadia subduction (Fig. 37, Phase 1). This deformation produced numerous north-trending folds, north-trending reverse and thrust faults, and west-northwest–trending steep reverse (?) faults within the Paleogene bedrock (Madin et al., 1995) (current orientations). This style of deformation is similar to active Cascadia fold and thrust belt deformation (Goldfinger et al., 1992).

From the South Jetty the opening of the old jetty railroad tunnel is visible (Fig. 48D). At the upper left corner of the tunnel opening and northwest across the cove cliff we can see a single water laid tuff ~30 cm (12 in.) thick (Fig. 48E). This tuff of Cascade arc magma type has been dated at 8.1 ± 0.1 Ma (M. Darin, personal commun., Oct. 2020; I.N. Bindeman, personal commun., April 2021 and June 2021).

Our next stop is within the community of Charleston. Drive back along Coos Head Road to the intersection with the Cape Arago Highway and turn left toward Charleston. In Charleston, before the bridge, turn left onto Boat Basin Road and drive northwest as far as permissible. The stop is in the cove at the end of the road, past the Coast Guard residence building and facilities of the Oregon Institute of Marine Biology (OIMB). The cove access is across private property and permission should be obtained from the Director of OIMB, P.O. Box 5389, Charleston, Oregon 97420; (541) 888-2581.

Upon arriving at the cove proceed down onto the old concrete jetty and walk to the outer end. From this vantage point, we can observe the Whisky Run Terrace truncating the Empire Formation sandstone. The Charleston fault trends through the gully that separates upthrown Empire on the far northwest side at an elevation of ~20 m (65 ft) in contrast to the southeast side at ~3 m (10 ft). Detailed mapping of the Whisky Run Terrace has demonstrated that many of these faults are “active” (McInelly and Kelsey, 1990; Madin et al., 1995).

The Whiskey Run Terrace deposits include littoral and aeolian sands with a deep soil horizon. The truncated surface of the Empire Formation is bored by the extant pholad Penitella (large conical borings typically into sandstone-eroded to shallow-bowls) and Adula (heart-shaped cross section, typically bored into mudstone or softened sandstone), part of the early “rock boring” community (Addicott, 1964; Evans, 1967; Armentrout, 1975; Hiebert, 2015). The boring cavities provide living space for other endolithic organisms, including nesting pelecypods, juvenile sea anemones and crabs.

***SPECIAL DRIVING INSTRUCTION FOLLOW for Stops 17, 18, and 19: the pull-out at the next stop is small and the pavement edge has a sharp drop-off to the gravel area!

Drive back through Charleston turning left (north) on Cape Arago Highway toward Coos Bay. Check mileage: From the center of the Charleston bridge across South Slough, drive 1.6 miles north just past the Barview Market store to the small left-side gravel pullout at a concrete block building slightly before Beacon Road on the right (Fig. 49). This informally named Beacon Lane Cove area provides access to the Empire Formation and the Coos Conglomerate at Fossil Point. The traverse is about a half mile and requires at least one hour’s time on the outcrop. The western point of this cove must have a broad exposure on a falling tide to assure you will have access to the Fossil Point and be able to safely return. All of the Whisky Run Terrace along this part of the Coos Bay Shipping Channel is private property. (Parking at 43.358206°N, 124.308641°W.)

WARNING ABOUT RISING TIDES: The western point at Beacon Lane Cove becomes a rising-tide barrier to dry exits from the Fossil Point area. Collectors become enraptured with the abundance and diversity of fossil within the Coos Conglomerate at Fossil Point and may miss a dry exit back to their vehicle.

This traverse is across the Empire Formation type section along the northeast limb of the South Slough Syncline. The type section is just over 700 m (2310 ft) thick and is thin bedded in core samples, but outcrop surfaces suggest extensive bioturbation with a mottled weathering patina. The Empire Formation unconformably overlies the early-middle Miocene Tarheel formation along the northeast limb of the South Slough Syncline (Fig. 20). The basal contact with the underlying Tarheel occurs on the sea floor surface adjacent to the type Tarheel outcrops north of Pigeon Point along the northeast limb of the South Slough Syncline (Figs. 19 and 20, and Stop 19, Fig. 50).

The molluscan fauna of the Empire Formation was originally studied by Dall (1909) who assigned it to the Pliocene Epoch. Biostratigraphic analysis of diatoms reassigned the fauna to the Wishkahan Molluscan Stage of the late Miocene (Addicott, 1976; Barron, 1981). Fossiliferous beds at Cape Blanco are considered coeval with the Coos Bay Empire (Addicott, 1980). The fauna is interpreted as being deposited in inner sublittoral environments (Armentrout, 1967).

The Empire Formation is fossiliferous with 114 molluscan species, most listed from the Coos Conglomerate, and rare but paleontologically important marine mammal fossils including whale, seal, sea lion, walrus, and Desmostylus (Dall, 1909; Ray, 1976; Roth, 1979). The Coos Head tuff dated at 8.1 ± 0.1 Ma is tentatively correlated with the interval of Beacon Lane Cove area but has not yet been observed due to sand and mud cover (D. Blackwell, pers. commun., May 2021).

The shore face traverse south from Beacon Lane Cove to Fossil Point (Fig. 49A and 49B) encounters numerous fossils of the large mollusk Patinopecten representing at least two species (Fig. 49C). Small bank outcrops beneath the Whisky Run Terrace expose excellent examples of the ubiquitous Empire trace fossils that include several centimeter-diameter vertical burrows and abundant lateral traces of similar width but that appear to be surface or very shallow trails. At very low tide concretionary curvilinear deformation patterns occur in the massive sandstone that suggest flexural slip related to the post-Miocene folding event of South Slough Syncline (R. Weldon, personal commun., June 2021) (Fig. 37, Phase 3).

Coos Conglomerate is exposed at “Fossil Point,” a prominent point along the shoreline traverse (Fig. 49B). Access onto the outcrop is by way of a bush-crawl trail off to the left of the photo Figure 49B (labeled on Fig. 49A), or visit at a significant minus tide. All adjacent property along this area of Whisky Run Terrace is private.

The spectacularly fossiliferous Coos Conglomerate (Fig. 49D) has been interpreted either as an intraformational member within the Empire Formation (Howe, 1922; Weaver, 1945) or as a younger deposit unconformably over the Empire Formation (Diller, 1899, 1901) (see also Armentrout, 1973). The conglomeratic section is 11 m (36 ft) thick and exposed across an area of less than 0.004 km² (less than one acre). The depositional environment interpreted ranges from a shoreface channel with longshore drift accumulation of the conglomerate (post Empire model: Diller, 1901) to a fluvially sourced channelized deposit contemporaneous with estuarine deposition (intraformational model: Weaver, 1945).

The base of the Coos Conglomerate is clearly erosional (Fig. 49B). The deposit consists of at least eight successions of vertically and laterally fining conglomerate. The clasts include rounded cobbles and boulders of Empire-like sandstone, some with weathering rinds, suggesting the rocks were resident in a vigorous weathering regime and then transported possibly in successive flood events.

The coarse sediment and “laterally shingled” fining cycles suggest a fluvial bar deposit sourced from an area where vigorous erosion of Empire strata resulted in concentration of fossil material. There are no reported fossils of early Miocene age from the highly fossiliferous and unconformably underlying Tarheel formation. The basal Coos Conglomerate can be traced ~30 m (100 ft) northwest fining until there are no pebbles or shell fragments, appearing to grade toward and possibly into the Empire Formation proper. However, the traceable contact is eroded out before a truly interbedded relationship can be confirmed. The top of the uppermost fossiliferous conglomerate appears to dip “beneath” Empire Formation sandstone, but persistent beach sand buries the “area of contact” precluding confirmation of an interbedded relationship.

Preliminary analysis of fracture patterns in the Empire-proper versus Coos Conglomerate suggests different deformational histories, but the friable-sandstone versus highly cemented conglomerate may account for the observed difference. Zircon age studies may provide addition criteria for the relative age of the two deposits. If the Coos Conglomerate is unconformably overlying the Empire Formation the unconformity may correlate with the late Miocene ca. 7.5–6 Ma global unconformity interpreted by McNeill et al. (2000) in the offshore Newport Basin (Fig. 37, Phase 3 deformation).

The Coos Conglomerate contains basaltic pebbles probably sourced from early Eocene igneous rocks to the southeast toward Roseburg. The Empire Formation sandstone proper is not known to contain these basaltic pebbles either below or above the Coos Conglomerate suggesting two possibilities: (1) the conglomerate includes in-part a different source area than the Empire Formation alone suggesting a post-Empire sediment provenance; or (2) the studied Empire exposures do not include the area where basaltic clasts were common, suggesting the coarse conglomerate came from an area of Empire deposition that is either unexposed or totally removed in the post-Empire deformation and erosion.

There is a second fossiliferous deposit higher in the Empire Formation type section that is clearly interbedded with the typical Empire sandstone facies. Weaver (1945, plate 9) shows this conglomerate as an offshore reef. Careful examination of the mapped location in outcrop suggests the placement on Weaver’s plate 9 is a drafting error and the upper conglomerate belongs within the shoreline section.

Return to your vehicle and continue the drive northeast for another 1.4 km (0.9 mi.) to the vicinity of Hedge Lane and Grinnell Lane, parking on the northwest side of the Cape Arago Highway (Fig. 50). The trail to the tide flat along the Coos Bay Shipping Channel is opposite Grinnell Lane. This locality requires a moderately low tide to access the seasonally exposed intertidal outcrops.

Arriving at the beach, note the ancient tree stumps similar to those along Big Creek at Sunset Bay. Just southwest of these rooted stumps are several seaweed- and barnacle-encrusted discontinuous concretionary horizons that are fossiliferous (Fig. 50A). This is the area of Tarheel formation strata, the “Miocene beds” (Tarheel formation of Armentrout, 1967) listed in the Charleston Quadrangle text (Madin et al., 1995). The concretionary strata occur in seasonally exposed fine-grained sandstone underlying the beach sand and bay mud.

In the 1940s, the Coos Bay shipping channel was dredged to 10 m (30 ft) depth from mile 2.0 to mile 4.5 and a Miocene molluscan fauna was recognized in the spoils (see Moore, 1963, Fig. 3). The dredged fossils were initially collected by Ellen James (Moore) for her University of Oregon master’s thesis. Dredging history is described and fossils are monographed along with the coeval Astoria Formation mollusks in Moore (1963). The most prevalent species recovered from the dredging spoils was and still is Dosinia whitneyi (Fig. 50D).

Outcrops along the east side of the shipping channel were discovered with a diagnostic early to middle Miocene molluscan fauna including Dosinia whitneyi, Chione ensifera, Crepidula praerupta and Psephaea indurata and informally named the Tarheel formation by Armentrout (1967, 1978, 1980). The lowermost Empire in this Barview section unconformably overlies the Tarheel with an angular discordance of ~12 degrees. Approximately 75 m (246 ft) of Tarheel strata are exposed with a dip of 35 degrees. The base is not exposed but projection of formation boundaries on the Charleston Quadrangle (Madin et al., 1995) suggest the possibility of a Tarheel thickness of 600 m (1970 ft) probably overlying with unconformity the Oligocene Bastendorff Shale. Occurrence of the Bastendorff Shale is based on recovery of Oligocene foraminifera at the northern end of the dredged spoils (Moore, 1963) (Fig. 50).

The Tarheel formation is restricted to the South Slough Syncline as is the Empire Formation. The molluscan faunas consist of forms living in shallow to moderate depths with a bottom sediment of fine-grained sand to silty mud. This suggests a moderately shallow estuarine embayment aligned along the northwest to southeast Neogene synclinal axis. The Tarheel/Empire unconformity is middle Miocene age, calibrated by zircon ages of 18.1 and 15.9 (see Fig. 16; also Fig. 37, Phase 2 deformation). This unconformity correlates with the Astoria and Columbia River Basalts interval at Yaquina Head near Newport, Oregon (Armentrout et al., 1983; Prothero et al., 2001a.

This Coos Bay area trip has traversed a lot of stratigraphy and addressed 56 million years of tectonic history. For the most part, the Coos Bay Cenozoic history matches the regional interpretations of Wells et al. (2014) for the Paleogene, and the interpretations of McNeill et al. (2000) for the Neogene. Zircon-age calibration of the Cenozoic formations in the southern Cascadia forearc basin result in affirming most previous biostratigraphic correlations. This suggests that much of the Coos Bay stratigraphic record represents what was differentially eroded off the Coast Range Crest and is really a western “remnant” to the Paleogene “Tyee-Coos Bay” basin, here-in referred to as the Paleogene Cascadia Basin. Zircon ages and biostratigraphic data also encourages the extension of the Paleogene Cascadia Basin westward beyond the Fulmar fault and offshore Pan American and Fulmar wells. The paleomagnetic rotation estimate of ~70° counterclockwise rotation suggests the Coos Bay area Eocene paleoshoreline was relatively north-south similar to the modern shoreline.

The one regionally significant “discovery” is the late Oligocene–early Miocene geographic isolation of the Coos Bay area from Cascade arc river sediments, suggesting early uplift of the southern Oregon Coast Range. This new interpretation merits continued testing with zircon population studies in other forearc basins and perhaps exhumation studies along the crest of the Coast Range.

Oregon’s coastal Highway 101 traverses a varied course across lowlands and river valleys as well as over several volcanic headlands before turning east on Highway 18 across the Coast Range and Tualatin Valley to Portland.

From Coos Bay north to Florence are 80 km (50 mi.) of coastal dunes, the Oregon Dunes National Recreation Area, the largest coastal dunes in North America. Reaching nearly 150 m (500 ft) in height the dunes enclose freshwater lakes and small areas of spruce and cedar forests. Campgrounds and hiking trails are abundant, as are dune buggy rides. Much of the sand has been reworked landward from lowstand shorelines (Peterson et al., 2007). Recent studies have determined that the youngest dunes, which were formed over the past seven thousand years, are nearest the ocean. The higher dunes to the east were formed more than 20,000 years ago, and the tops of some of the higher dunes were last active more than 100,000 years ago. Analyses of the chemical makeup of individual sand grains point to the Umpqua River, just west of Reedsport, as one primary source of the Oregon Dunes sand, with contributions from the Siuslaw and other, smaller rivers (Cloyd, 2018). Peterson et al. (2007) studied dunes along the entire Oregon Coast and interpreted a much more complex history with longshore transport from the south during lowstand as a primary sediment source. Lowstand accumulations were reworked landward by landward waves as rising Holocene sea level transgressed the forearc margin.

Oregon’s coastal elevation varies not only from rock type but also varying rates of uplift resulting from intrinsic strain accumulating along the Cascadia subduction boundary (Fig. 51). Based on geodetic data comparing 1980s measurements with those from 1930 to 1941 the high areas are interpreted as evidence of intrinsic elastic (temporary) strain causing a bulge between the Juan de Fuca and North American plates (Figure from Orr and Orr, 2012; constructed from original work cited there in). Subduction earthquakes of 8+ magnitude occurring every 300–500 years may release this strain. A new early warning system is now in place to provide several minutes of warning of earthquakes for critical infrastructure and citizens to prepare for the shaking.

North of the dune area, Highway 101 begins traversing a series of headlands of late Eocene Yachats Basalt (36–34 Ma). The Yachats basalt and farther north basalt headlands at Cascade Head and Tillamook Head (Tillamook episode) are the late phase of Yellowstone magmatism erupted through the accreted Siletzia terrane (Wells et al., 2014; Camp and Wells, 2021). These eruptions mark the beginning of Eocene regional margin-parallel extension and widespread dike injection in the Cascadia forearc.

Cape Perpetua north of Florence and Cascade Head north of Lincoln City are composed of these tholeiitic basalts. The Cape Perpetua Special Interest Area provides access from ocean side parking along Highway 101 to paved walkways through Yachats basalts and volcanic conglomerates cut by dikes. Walking along these trails and viewing the exposure of the many flows exposed in the cliff of Cape Perpetua provides an appreciation of this phase of late Eocene extensional volcanism of the forearc (Davis et al., 1974; Wells et al., 2014; Camp and Wells, 2021). The narrow, winding drive to the top of Cape Perpetua affords a spectacular view of Oregon’s rugged coastline. The road to the top is marked for the Cape Perpetua Campground and is north of the exit to the Cape Perpetua Visitor Center, another worthwhile stop.

Farther north at Newport and Depoe Bay are exposures of Columbia River Basalt Group flows sourced from fissures in eastern Oregon and Washington 15–12 million years ago (Beeson et al., 1989; Wells et al., 1989) (Fig. 4). Magnetostratigraphic and geochemical studies have “finger-printed” many of the Columbia River Basalt Group flows which number in the 80s. Yaquina Head north of Newport, the harbor and headlands at Depoe Bay, and Cape Foulweather are flows of the Columbia River Basalt Group. Depoe Bay exposures below the sea wall provide views of pillow basalts, a “spouting horn,” and often, farther offshore, whales (Reidel et al., 2013; Miller, 2014).

Some Columbia River Basalt Group lava flows, having high bulk density, pooled in the less dense water-saturated sediments upon reaching the coast, deforming the sediments plastically, and intruding the brittle older strata forming invasive dikes and sills, for example the ring-dikes visible from Cape Foulweather (Niem et al., 1994). The Cape Foulweather basalts have been correlated with the Gingko Flow (15.6 Ma) in central Washington east of the Cascade Mountains (Wells et al., 1989). Offshore drilling encountered Grande Ronde Basalt at depths below 1070 m (3500 ft) and a sill at 2500 m (8250 ft) (Niem et al., 1994; Wells et al., 2009).

The 56–49 Ma basalts of the Siletzia oceanic plateau are not encountered on our drive but core the Coast Range east of the coastal highway. Comparison of Siletzia with the Hawaiian Islands gives a sense of the scale of the Coast Range portion of the accreted terrane (Fig. 52).

From Lincoln City, Highway 18 takes us across the Coast Range to the agricultural area of the Tualatin Valley, one of Oregon’s premier wine growing areas. Of note along this route west of McMinnville, is Erratic Rock State Natural Area, a Proterozoic Belt/Purcell argillite block ice-rafted from Montana or southern Canada on one of the Missoula Floods ~15,000 years ago (Fig. 3). Mapping glacial erratics throughout the Willamette and Tualatin Valley defines the shoreline of the floods (Allison, 1935, Bretz, 1969; Waitt, 1985; Minervini et al., 2003).

Field partners David Blackwell, Laird Thompson, Christine Rossen, Noel Blackwell, and Rebecca Dorsey have contributed their insightful field observations reflected throughout this field-trip guide. Additionally, Tom Armentrout and Jeffrey Armentrout provided field support on several traverses.

Portions of this text have been read by and improved with edits from David Blackwell, Michael Darin, Jim Jackson, Les Magoon, William Orr, and Laird Thompson. GSA Field Trip Chair Adam Booth and GSA Publications staff edited the entire manuscript and significantly clarified several elements. This final version is John Armentrout’s distillation of observations, data, and discussions with the entire Coaledo Project Team.

Coaledo Team members and associates include John M. Armentrout, Mobil Oil, retired: project coordinator and stratigraphy; Allison Barbato, Louisiana State University: palynology and geochemistry; Ilya Bindeman, University of Oregon: tuff geochemistry; David Blackwell, University of Oregon: stratigraphy/structure/paleomagnetism; Noel Blackwell, University of Oregon: paleomagnetism; Scott Bogue, Occidental College: paleomagnetism (AMS); Sam S. Cooke, University of Oregon: petrography; Michael Darin, University of Nevada-Reno: zircon analysis and regional tectonics; Rebecca Dorsey, University of Oregon: sedimentology and regional tectonics; Thomas Demchuk, Louisiana State University: organofacies and maturation history; Carole Hickman, University of California at Berkeley: molluscan paleontology; Jim Jackson, Portland State University: structure/tectonics; Rocky Johnston, PacWest Drone Services: drone pilot; Susan Kidwell, University of Chicago: sedimentology: Leslie Magoon, U.S. Geological Survey, retired/Stanford University: petroleum systems; Kristin McDougall, U.S. Geological Survey: micropaleontology; Christine Rossen, ExxonMobil, retired/consultant: sedimentology; Peter Ruggiero, Oregon State University Oceanography: wave modeling; Laird Thompson, Mobil Oil, retired: structure/tectonics; Dean Walton, University of Oregon: drone pilot; Sophie Warny, Louisiana State University: palynology; Ray Weldon, University of Oregon: paleomagnetism; Bruce Welton, New Mexico Museum of Natural History: elasmobranchs.