Santa Cruz Island field trip: Geology, history, and research opportunities
Published:May 18, 2020
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Thomas L. Davis*, Richard J. Behl†, Katie M. O’Sullivan†, Sarah Raskin, Stephen Bryne, 2020. "Santa Cruz Island field trip: Geology, history, and research opportunities", From the Islands to the Mountains: A 2020 View of Geologic Excursions in Southern California, Richard V. Heermance, Joshua J. Schwartz
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*Principal organizer and corresponding author: email@example.com.
This field trip provides a rare opportunity to visit outcrops and structures that highlight the geology, history, and natural beauty of Santa Cruz Island, a remnant of prehistoric California isolated off Santa Barbara. Santa Cruz Island provides some of the most southwestward positioned subaerial outcrops of the North American landmass, while displaying a rare glimpse of prehistoric coastal southern California and picturesque and seldom accessible exposures of Tertiary strata. Most of the stops are difficult to reach and many are closed to public access. Stops 1, 9, 9B, 9C, 13, and 13B are within the Channel Islands National Park, and access to the park portion of the island is by public boat transport (Island Packers) from Ventura Harbor to Prisoners Harbor. Stop 1 is near the pier at Prisoners Harbor and easily accessible; however, the other stops require roundtrip hikes of at least 10 miles from the pier. One of the goals of this four-day trip is to visit as much of the island’s varied geology as possible. A significant body of widely recognized geologic research has been done on the island from late Quaternary sea-level and climate changes to the tectonic evolution of the western North American plate boundary, and in particular the transformation of a subduction to transform plate boundary along a continental margin. Discovery that SCI and the western Transverse Ranges have rotated ~90° clockwise since the early Miocene (Kamerling and Luyendyk, 1979, 1985; Luyendyk et al., 1980) brought on an intense period of research on the island from the late 1970s through the 1990s. Much of this work has been published in both the formal and informal literature. Two decades later, this field trip is an opportunity to review much of these additions to geologic understanding with the advantage of gains in knowledge since then. The guide will emphasize each stop’s importance, offer questions for future research, and showcase the island’s earth science educational opportunities. This four-day trip requires 4WD vehicles and includes some 3–6 km (~2–4 mile) hikes.
Dedicated to Dr. Lyndal Laughrin, Santa Cruz Island Reserve Director, Emeritus, The Sage of Santa Cruz Island
As mentioned before, one of the goals of the trip is to visit as much of Santa Cruz Island’s (SCI) varied geology as possible in four days. Field-trip days are valuable, given the logistical difficulty to reach many of the stops, and for many even reaching the island. There will be more opportunities to read, review and debate the geology of the island in the evening at the University of California Field Station, or back on the mainland. Accordingly, the field-trip leaders have designed the Road Log, Topical Notes, and figures to emphasize key observations at each stop. Many of the stops refer to Topical Notes that follow the Road Log and place the stops into a larger geologic context. Figures 1–5 serve to introduce the island’s geologic framework and field geology, and it is advised that trip participants become acquainted with these figures initially: Figure 1 shows the trip route and stops, Figure 2 is a composite satellite image with stops, Figure 3A is a generalized geologic map with stops, Figure 3B has cross sections, and Figures 4 and 5 are generalized stratigraphic columns. This field guide contains more than geology, though. SCI’s late Pleistocene and Holocene isolation from the mainland due to sea-level rise has provided the island with unique flora and fauna that Sarah Raskin summarizes in the “Flora and Fauna of Santa Cruz Island” section. Humans arrived on Santa Cruz Island during the latest Pleistocene and made the island home until the 1820s when the Chumash people left the island (Laughrin, 2015). In the section on “Prehistoric Chert Use on East Santa Cruz Island,” Stephen Bryne outlines the indigenous people’s use of the island’s chert resources, including the “Santa Cruz Island Blonde chert,” for a variety of tools that were used on the island, on other Channel Islands, and across southern California. Chert may be seen in Stops 9B, 9C, and 13B.
On the first day of the trip, there will be a geologic orientation at Prisoners Harbor, a drive along the Canada del Puerto to the University of California Santa Cruz Island Field Station, check-in at the station with a facility orientation, and lunch at the station, followed by a driving reconnaissance of the central portion of the island (Fig. 1). On the second day, we will return to Canada del Puerto to look at the Santa Cruz Island (SCI) Volcanics, and then drive eastward along Navy Road, which follows the crest of the island’s eastern isthmus. Then we descend the Chinese Harbor Road to Monterey Formation outcrops where we make two hikes: one to a burnt shale outcrop and the other to a map-scale anticline. Lunch will be at Chinese Harbor. On the third day we will drive along the crest of the Southern Ridge to the western portion of the island, viewing drainage offsets along the SCI fault and the Willows Plutonic Complex along the way. We will then head across the southwestern portion of the island to Poza Canyon. There we will make a hike to the mid-Tertiary beach section between the mouth of Poza Canyon and Near Point (lunch is at the beach). Returning to the vehicle, we will then drive to Christi Beach and make two hikes: one to the lower and mid-Tertiary outcrops south of Christi Beach and then a hike to a sea-cliff exposure of the Santa Cruz Island fault. On the fourth and final day, we will pack our gear, help clean the Field Station, and drive through the Canada Puerto and eastward along the Navy Road. At the end of the Navy Road, we will make a hike to the contact between the Monterey Formation and underlying Santa Cruz Island Volcanics along the western edge of Montañon Ridge—aka El Montañon highlands (see section on “Prehistoric Chert Use on East Santa Cruz Island”). We will have lunch here and drive to Prisoners Harbor for the boat return to Ventura Harbor. Be sure to occasionally refer to the “Flora and Fauna” section during the four days of driving and hiking SCI.
For those wanting more field information, there are several very good guidebooks and geologic maps about the island. Guidebooks by Weaver et al. (1969), Boles and Landry (1997), and Boles (2015) have an abundance of field descriptions and sketches, maps, cross sections, stratigraphic columns, photos, and general background. Unfortunately, most of these materials are out of print, but an online search may make locating copies possible. The geologic maps of Weaver and Nolf (1969) and Dibblee (2001a, 2001b) are equally important. The volume by Weaver et al. (1969) that contains the Weaver and Nolf map is out of print. Fortunately, the U.S. Geological Survey (USGS) and National Park Service (NPS) have put the Weaver and Nolf map into GIS format and it is available online for free download. The map is titled “Digital Geologic-GIS Map of Santa Cruz Island, California (NPS, GRD, GRI, CHIS, SCIS digital map) adapted from an American Association of Petroleum Geologists Field Trip Guidebook map by the University of California, Santa Barbara Geological Survey (1969)” and can be found at https://catalog.data.gov/dataset/digital-geologic-gis-map-of-santa-cruz-island-california-nps-grd-gri-chis-scis-digital-map-1969. The Dibblee maps are equally important and are available for purchase from the Santa Barbara Museum of Natural History, Museum Store, https://www.sbnature.org/store/.
A NOTE ABOUT VISITING THE ISLAND AND SAFETY
If you are visiting the Santa Cruz Island Reserve without the GSA 2020 Cordilleran Section Meeting field-trip leaders, you must fill out the online application (https://rams.ucnrs.org/) and make a separate reservation for each visit. Please note that the reserve is only open to researchers and educational organizations; the reserve is not open to the public. The Reserve Director is in charge of approving requests. After you have received approval to visit, you must reserve transportation to the island (https://santacruz.nrs.ucsb.edu/visiting/transportation).
Don’t forget to bring a flashlight! There are no stores on the island and cell coverage is spotty. The Field Station recommends bringing an extra day’s supply of food in case of an unforeseen emergency and participants are responsible for their emergency supply. Individual cooking on camp stoves is strictly prohibited, due to possible fire hazard. Read the Safety page (https://santacruz.nrs.ucsb.edu/visiting/safety) for “What not to bring” to Santa Cruz Island. Transportation on the island will be in 4 WD vehicles rented from the Field Station.
Day 1. Friday, 15 May 2020, 0.4 ft low tide @ 12:06 p.m. PDT
We will have an orientation near the pier to discuss safety, itinerary, objectives of trip, guidebook, and to provide an overview with geologic maps.
Stop 1. Prisoners Harbor (34.019408°N, 119.684287°W)
Santa Cruz Island is located along the southwestern edge of the western Transverse Ranges segment (Dibblee, 1982a) and adjacent to the Continental Borderland Province (Crouch, 1979, 1981; ten Brink, 2000; Marsaglia et al., 2019) that is located just to the south of the island (Fig. 2). The surface geology of the island has structural and lithologic similarities to the mainland part of the western Transverse Range segment: a large east-west–trending anticline, a large left-lateral, strike-slip fault (Santa Cruz Island fault), a thick sequence of Cenozoic sedimentary and volcanic rocks, and oceanic-affinity basement rocks (Dibblee, 1982a, 1982b).
Santa Cruz Island is the largest of the Channel Islands of southern California. In the most general sense, SCI is an antiform composed of three anticlines separated by two structurally complex synclines. The strong surface expression of these folds is a result of differential erosion of the various rock units and most likely young crustal convergence (Figs. 2, 3A, 3B). Prisoners Harbor is on the north limb of the central anticline, and along the west side of the harbor are north-east dipping layers of the Prisoners Harbor member of the SCI Volcanics (Nolf and Nolf, 1969). The central anticline is east-west–trending and expressed at the surface by two east-west–trending ridges separated by a long central valley (Fig. 1). The Northern Ridge is underlain by north-dipping and resistant Miocene SCI Volcanics, the central valley is developed over weakened rocks along the SCI fault and its branches and splays, and the highest portion of the Southern Ridge is underlain by Mesozoic crystalline rocks of the SCI Schist and Willows Plutonic Complex, with the lower ridge flanks underlain by south-dipping Miocene volcaniclastic and other sedimentary rocks. Cross sections of the central anticline and its relationship to the Northern Channel Islands anticlinorium are shown in Figure 3B (Shaw and Suppe, 1994; Pinter et al., 2003). The northwest-trending Christi anticline, in the southwest of the island, exposes lower to middle Tertiary strata. A narrow isthmus of landslide-prone Monterey Formation lies in a structural trough or synclinorium that separates the central and northeast anticlines. The northeast anticline trends northwest, and its steep southwest limb underlies Montañon Ridge, the dominant topographic feature of this part of the island. It’s noteworthy that northwest structural trends are subparallel to the borderland province structures that lie to the south. The northeast anticline is underlain by SCI Volcanics covered with smaller patches of Monterey Formation. Some of these patches appear to be small half-grabens with Monterey Formation preserved in the hanging wall. The anticline is broad-crested, faulted, and its northeast limb lies offshore (Redin et al., 1999). Almost no late Miocene or Pliocene rocks are preserved on Santa Cruz Island. Cross sections of the nearby offshore show that rocks of this age may have been removed by subsequent erosion related to the initiation and uplift of the modern Transverse Ranges. (See topical note section on the “Transverse Ranges.”)
Depart Stop 1 and drive along the Canada del Puerto to the central valley. The route is southwesterly and takes us down-section through the four members of the SCI Volcanics sequence. From youngest to oldest, the members are the Prisoners Harbor, Devils Peak, Stanton Ranch, and Griffith Canyon (Fig. 4). The volcanic sequence consists mostly of flows, flow breccias, and dikes. Compositionally, the sequence is basaltic andesite, andesite, and dacitic andesite. The sequence is up to 2000 m thick and early to middle Miocene in age (17–16.3 Ma; Luyendyk et al., 1998). During Day 2, we will return to the Canada del Puerto and make two stops (7 and 8) to examine the SCI Volcanics in more detail. After several kilometers, the Canada del Puerto opens into the east-west–trending Canada del Medio (central valley), and our route swings to the right and follows the north strand of the SCI fault for a short distance along the north edge of the central valley. Our route crosses the broad wash and swings to the left, and the UC Field Station is on the left.
Stop 2. UC Field Station (33.996897°N, 119.726170°W)
Station orientation by facility director and staff. Check into rooms, prepare for the field, and have lunch. Depart the station, head west a short distance along the poorly exposed trace of the south strand of the SCI fault, turn left, and climb the road to the crest of the Southern Ridge (Fig. 1). The road climbs through outcrops of the SCI Schist that weathers to a dark-red soil. At the crest, take the left branch of the Southern Ridge Road (see “Flora and Fauna” section).
Stop 3. Southern Ridge above Willows Canyon (33.996650°N, 119.753896°W)
From here, much of the south side of the island is visible, and the soil colors are helpful for distinguishing some of the principal rock units and structure of the central anticline (Figs. 1–3, 5). Looking south into Willows Canyon, the brick-red soil marks the extent of the SCI Schist and, further down canyon, the whitish-gray soil is above the Willows Plutonic Complex. Further down the canyon are south-dipping, light-tan and whitish-gray strata of the early to middle Miocene Blanca Formation that consists of thick-bedded volcaniclastic conglomerate, tuffaceous sandstone, and tuff. The gray to gray-blue low ridges along lowermost Willows Canyon, near the coast, are Lower Miocene San Onofre Breccia (Fig. 6) that is uplifted along the south-dipping, east-west–trending Willows fault. Stop 3 is along the south-dipping limb of the central anticline (Shaw and Suppe, 1994; Pinter et al., 2003) that has been interpreted to be the result of late Pliocene and Quaternary north-south–shortening associated with development of the western Transverse Ranges (Fig. 3B), or Miocene extension that resulted in hanging wall block rotation (Baker, 1993), or some combination of both events. (See topical note on the “Central Anticline.”)
To the north is the steep drop into the Canada del Medio with the trace of the north branch of the SCI fault at the break in slope along the north edge of the valley. This view is dominated by the rugged Northern Ridge that is underlain by SCI Volcanics, and to the left is the highest point on the island, Diablo Peak (740 m, 2429 ft). At this stop, we will take a closer look at the SCI Schist. As described by Boles (2015), the SCI Schist is an olive to gray-green chlorite schist that weathers to a brick-red color. The schist is fine grained and well foliated with minor greenstone and milky quartz veins. The schist is the oldest rock on SCI, and the protolith is intermediate-composition volcanic rocks and sediments (Hill, 1976). Its age is uncertain but possibly Jurassic. Willows Plutonic Complex (162 ± 3 Ma; Mattinson and Hill, 1976) and the Alamos Tonalite (141 ± 3 Ma U/Pb; Mattinson and Hill, 1976) intrude the schist. A 130–120 Ma K/Ar age of the schist suggests an early Cretaceous thermal event (Mattinson and Hill, 1976). (See topical note section on “Exotic Terranes.”) Depart Stop 3 and drive southeast along the Southern Ridge Road, which is underlain by brick-red soil and the schist, to Stop 4 near Hill 1484.
Stop 4. Near Hill 1484 (33.986651°N, 119.739492°W)
This stop is along the contact between the middle member of the Blanca Formation and the underlying SCI Schist. Here the contact is easy to reach but somewhat poorly exposed. Mapping shows that the lower Blanca member and basement rocks, present to the west, are missing here (Fig. 7). Weaver and Nolf (1969) show the contact as faulted (Anchorage fault), while Dibblee (2001a) shows the contact as an unconformity that is overlapping the older units. If it is a fault contact, then the map pattern indicates it is low angle and south dipping, and the missing older units raise the question of whether it is a significant denudation fault associated with Miocene extension (Baker, 1993), or merely a faulted unconformity. If the contact is an unconformity, then its map pattern and missing lower units show the Blanca Formation was deposited across a surface with local topography. Further west, a breccia containing Willows Plutonic Complex clasts in the lower Blanca Formation supports some local topographic relief (Boles, 2015), and argues against a significant denudation fault. Before leaving Stop 4, we will examine the lithology of the Blanca Formation just above its basal contact with the schist. The Blanca consists of volcaniclastics and pyroclastics with interbedded sandstones (McLean et al., 1976). As with the San Onofre Breccia, the lower members of the Blanca contain pieces of blueschist. Geochemical data suggest that the clasts within the Blanca, while volcanic in origin, do not come from the SCI Volcanics, Conejo Volcanics, Zuma Volcanics, or the Tranquillion Volcanics (e.g., Weigand, 1997). While the clasts within the Blanca do not appear to have traveled far from their source, the edifice remains elusive. Similarly, the age of the Blanca Formation is poorly constrained. A clast from the middle member has a K/Ar age of 13.3 ± 0.8 Ma, and a flow from the upper member has an age of 14.9 ± 0.8 Ma (McLean et al., 1976).
Turn the vehicles around and drive west along the crest of Southern Ridge Road and past Stop 3 to the panoramic view above Laguna Canyon that is Stop 5. Along the way, the road remains in the schist, and roadside outcrops display its foliated character and numerous quartz veins.
Stop 5. Southern Ridge above Laguna Canyon (34.007570°N, 119.786920°W)
Visible to the southwest is the prominent Sierra Blanca and its nearly complete section of south-dipping Blanca Formation (Fig. 8). East of the Sierra Blanca, folds and the Willows fault within the Blanca Formation are east-west trending, but the structural trend becomes northwesterly to the west of the Sierra Blanca. From this stop, we will briefly review the role of SCI in ideas about the tectonic history of southern California and plate margin tectonics. (See topical note section on “Santa Cruz Island and the Progression of Plate Margin Ideas.”)
To the north, beyond the central valley, is the Northern Ridge underlain by SCI Volcanics with Diablo Peak (aka Devils Peak) just east of due north. The eruptive center of the SCI Volcanics is postulated to be just east of Diablo Peak (Boles, 2015) and north of Canada del Portezuela (Nolf and Nolf, 1969).
Make a right turn off the Southern Ridge Road and descend a short distance to the saddle and drainage divide (Centinela) between the east-flowing Canada del Medio drainage and the west-flowing Canada Cervada drainage. The SCI fault cuts through the saddle and its shear zone is partially visible in local outcrop. Make another right onto the Canada del Medio Road and go east. Outcrops along this route show the road crosses the SCI fault several times with SCI Volcanics to the left (north) and SCI Schist to the right (south) (see the section on “Flora and Fauna”).
Stop 6. Portezuela (34.006867°N, 119.755719°W)
Stop 6 provides a good view eastward of the Canada del Medio (Central Valley) as it trends toward Valley Anchorage. Down-drainage of Stop 6 (to the east), the SCI fault splits into northern and southern branches that are positioned along their respective valley edges, showing the influence of the SCI fault on formation of the valley. In the far distance and left of the valley, the U.S. Navy facility on the crest of the isthmus portion of the island is visible. The isthmus is underlain by Monterey Formation and will be visited on Day 2 of this trip.
Depart Stop 6 and continue eastward along the Canada del Puerto Road toward the UC Field Station. If time allows, we will make two short stops along the road to look at the Griffith Canyon member of the SCI Volcanics (left of road and near small reservoir) and Blanca Formation outcrops (along the south side of road near crossing of the broad wash).
Day 2. Saturday, 16 May 2020, 0.6 feet low tide @ 12:49 p.m. PDT
Prepare for the field, have breakfast, and depart the UC Field Station at 8:00 a.m. Return to Canada del Puerto Road via the Canada del Medio Road connector (Fig. 1). We will drive toward Prisoners Harbor, making two stops along the Canada del Puerto Road to look at members of the northeast-dipping SCI Volcanics.
Stop 7. Santa Cruz Island Volcanics, Stanton Ranch Member (34.003370°N, 119.707800°W)
The outcrop is set of small, dark cliffs at the base of the hillside that are just to the north of the road and across the wash. The SCI Volcanics are divided into four members with overlapping ages, from bottom to top: Griffith Canyon, Stanton Ranch, Devils Peak, and Prisoners Harbor (Fig. 4). The Stanton Ranch Member lies conformably on the Griffith Canyon member (16.9 ± 0.2 Ma; Luyendyk et al., 1998) and is the second oldest member of the formation. The Stanton Ranch member consists of andesitic flows, breccias, and tuff breccias of varying thicknesses, and is thought to have been emplaced on an eruptive flank. Porphyritic andesite flows and hydrothermal alteration can be observed at this outcrop. Some of the phenocrysts within the porphyritic andesite are actually xenoliths of plutonic rock (Fig. 9), and, in rarer cases, pieces of glaucophane schist (Nolf and Nolf, 1969). There are several sets of slickensides visible from this stop (Figs. 10 and 11), most notably on the smooth surface ~5 m (~15 ft) above ground level (highlighted with white arrow in Fig. 10). Another set of slickensides indicates the presence of a west-dipping strike-slip fault that places andesite next to dacite (Fig. 10). (See topical note on “Santa Cruz Island Volcanics.”)
Stop 8. Santa Cruz Island Volcanics, Devils Peak Member (34.011880°N, 119.691880°W)
The outcrop is just to the north of the road along the far side of the wash. The Devils Peak Member of the SCI Volcanics consists of scoraceous andesitic and dacitic flows, flow breccias, and pyroclastics. Fresh surfaces are light gray and show abundant vesicles, while the weathered surfaces have intense staining that ranges from red to dark gray (Fig. 12). There is evidence that these flows were deposited on slopes adjacent to a volcanic edifice, which may be near the actual Diablo Peak; Fig. 1). The Devils Peak Member overlies the Stanton Ranch Member, with localized angular discordance where various amounts of reworking took place. The upper part of the Devils Peak Member was originally determined to have a K/Ar age of 16.1 ± 0.9 Ma (Turner, 1970), while a dike cross-cutting the lower parts of the member has ages of 16.0 ± 0.7 Ma and 19.9 ± 0.9 Ma (Crowe et al., 1976). More recently, using the Ar/Ar method, Luyendyk et al. (1998) determined the age of the lower part of the Devils Peak member to be 17.0 ± 0.1 Ma, while the upper part of the member was 16.9 ± 0.2 Ma. This discrepancy in numerical ages highlights the need for ages to be reexamined with additional sampling and more up-to-date techniques. (See topical note on “Santa Cruz Island Volcanics.”)
Depart Stop 8 and continue down the road to Prisoners Harbor and make a right on Navy Road. The lower part of Navy Road climbs through a highly weathered and deformed portion of the Prisoners Harbor Member of the SCI Volcanics. When the road reaches ~152 m (500 ft) elevation, outcrops of weathered and deformed Monterey Formation appear. Here, the basal contact of the Monterey Formation is poorly exposed and appears to be faulted. The road continues up the ridge through folded Monterey Formation, with the exception of a small outcrop of sheared and highly weathered SCI Volcanics near the crest. Navy Road turns eastward at the crest, continues through the isthmus portion of the island, and heads toward the Montañon Ridge and the northeast part of the island. Along the isthmus, the Monterey Formation is disrupted by numerous recent landslide masses and blocks. Closed topographic depressions are common and probably the result of down-slope block rotation closing off drainage patterns. Pass the U.S. Navy facility on the right and continue eastward about a half a mile. Turn left (north) on a smaller dirt road that descends steeply toward Chinese Harbor. On the fourth day of the trip, we will return to Navy Road and continue eastward along the crest of the isthmus to the end of Navy Road to examine the contact between the Monterey Formation and the SCI Volcanics. As we descend into Chinese Harbor, note the extensive flat surface to the left (west), at ~244 m (800 ft) elevation, that may be a landslide bench (Scott Minor, 2020, personal commun.) or the remnant of an uplifted Quaternary marine terrace. Weaver (1969) states that possible marine terraces occur as high as ~550 m (~1800 ft) above sea level on the north side of SCI.
Stop 9. Parking at Chinese Harbor (34.016051°N, 119.619563°W)
Park at a flat area at the toe of a recent landslide. To the east are whitish sea-cliff exposures of Monterey Formation and beyond the reddish-brown exposures of the SCI Volcanics at Coches Point, which is the northwest end of the Montañon Ridge—aka El Montañon highlands (Figs. 3 and 13). The view is not ideal for appreciating the map-scale folding at Chinese Harbor, but it is possible from here to see the southwest-dipping Monterey Formation and SCI Volcanics contact that are the southwest limb of the anticlinal structure that underlies the northeastern portion of the island. From the parking area, we will make a one-mile roundtrip hike to the burnt shale outcrops at the base of the sea cliff just northwest of the parking spot. After that, and dependent on the swell and tide conditions, we may make a 2–3-mile roundtrip hike to a map-scale anticline in the Monterey Formation. Both hikes are on cobble beaches that are best hiked with boots with ankle support. Lunch will be at the parking stop. (See the topical note on the “Monterey Formation of the Santa Barbara Channel.”)
Stop 9B. Chinese Harbor Burnt Shale (34.016111°N, 119.619917°W)
It is a short walk to the outcrop. We will examine an exposure of “burnt shale” in the Monterey Formation in the beach cliff at the seaward edge of the landslide that we drove over to reach Chinese Harbor. In contrast with the typically light-colored exposures of porous siliceous Monterey strata, this outcrop displays a striking pattern of deep-red, yellow, white, and black bands and seams that are unrelated to primary bedding (Figs. 14A–14B and 15). Burnt shale deposits are a global phenomenon occurring in organic-rich or bituminous mudstone that has undergone natural heating and combustion to various degrees, frequently resulting in near-surface metamorphism, mineralogic recrystallization, and even partial melting, forming “clinker rock” (Figs. 14C–14G and 15). Preliminary work on this exposure by Bedig Charkhutian and Kenton Crabtree (California State University Long Beach) has found the presence of high-temperature cristobalite and tridymite, quartz, plagioclase, and hematite instead of the clay minerals, opal-A or opal-CT, calcite, and pyrite typical of the siliceous organic shale protolith. Although many hypotheses have been proposed over the years (e.g., wildfires, lightning strikes, etc.), a common association with landslides was noted by Weaver and Meyer (1969, p. 97):
Occasionally a small area of dark brown, highly bituminous shale occurs near the base of the section as at the southern extremity of Yellow Banks Anchorage on the southeast coast of the island. There are numerous small areas, always associated with landslides, in which the shale has been burned and has thus acquired a reddish color, as for example toward the northeast end of Chinese Harbor, on the north side of the isthmus, where smoke is issuing from a recent landslide and noxious odors are emitted.
An explanation for this important relationship was refined from investigation of two modern “burning shale” events—one occurring in a landslide in the Los Padres National Forest (Mariner et al., 2008) and the other in the organic-rich La Vida Shale Member of the Puente Formation, Orange County (Boles et al., 2010). The latter was observed to become spontaneously heated, undergo near-surface natural pyrolysis, combust, and reach temperatures of >800 °C. Deposition of organic-rich sediment beneath anoxic seawater produces a highly reduced porewater environment in which the FeS2 polymorphs pyrite and marcasite can form when in the presence of iron. Oxidation of the marcasite, in particular, is a highly exothermic reaction. When carbonaceous sedimentary rocks bearing marcasite are normally weathered, the sulfides are slowly oxidized as atmospheric oxygen diffuses downward from the surface, perhaps mixed into groundwater. However, when a landslide or large rockfall (or excavation) instantaneously exposes a large amount of fine marcasite to a flush of oxygen, the exothermic reaction can produce enough heat to “crack” the kerogen, producing hydrocarbon liquids and gases that perpetuate and expand the combustion. The importance of sulfide oxidation in producing heat was recognized fully a century ago when deep boreholes were drilled to dynamite the major cut of the Panama Canal and the dynamite would prematurely explode from the heat in the boreholes released by the sulfide oxidation (MacDonald, 1920).
Similar burnt shale exposures to the one we will visit can be observed in many places in southern California, including in the Casmalia Hills, Orcutt field, Grimes Canyon, Point Dume, U.S. Highway 101 roadcut near Rincon Point, and Palos Verdes.
Stop 9C. Chinese Harbor Anticline (34.024422°N, 119.608950°W)
This stop involves a 2–3-mile roundtrip hike along cobble-covered beach to the anticline. We will view and discuss outcrop-scale structure along the west limb and hinge area of the Chinese Harbor anticline. Structures are developed in the well-laminated and thin-bedded Monterey Formation, and include parasitic second order “S & Z” folds, fault-bend and fault-propagation folds, ductile flow structures in fold hinge zone, and boudinage (Figs. 16–22).
The Monterey Formation in this area is composed of thin- to medium-bedded porcelanite, calcareous mudstone (shale), and limestone or dolomite. Originally diatomaceous, the siliceous sediments were altered by the increased temperature of burial to convert from biogenic opal-A to diagenetic opal-CT, making the rocks harder and less porous. Our view into the beach cliffs is chiefly looking along strike, and exposes a pair of map-scale NW-SE–trending folds—the Chinese Harbor anticline (Figs. 17 and 18) and a parallel syncline to the northeast. The interbedded lithologies have a tremendous contrast in strength (Fig. 22), and the folding is accomplished largely by bedding-parallel slip between layers of different rheology. As early as 1933, Rand described and interpreted progressive changes in tightness/openness of the folds in the Monterey Formation to be related to flexural slip and bedding-parallel detachments at the Chinese Harbor syncline. We agree and see this style of deformation whenever the thin-bedded siliceous strata of the Monterey Formation are in the diagenetic stage (opal-CT or quartz-phase). Flexural slip is also shown by numerous S- and Z- parasitic folds on alternate flanks of the anticlinal and synclinal axes (Fig. 19). Shortening is also accommodated by thrust faults within thicker beds (primarily carbonates), and reverse faults that step upward through packages of interbedded thinner porcelanite and shale beds. (See “Flora and Fauna” section and section on “Prehistoric Chert Use on East Santa Cruz Island.”)
After visiting Chinese Harbor, retrace route back to UC Field Station via Navy Road and Canada del Puerto.
Day 3. Sunday, 17 May 2020, 0.8 ft low tide @ 1:23 p.m. PDT
Prepare for the field and have breakfast. Day 3 involves a long drive to the western portion of the island and several stops and hikes. We will depart the UC Field Station at 7:30 a.m., head west a short distance, turn left and climb the road to the crest of the Southern Ridge, turn right and proceed about 6 miles, following a portion of the route we did on Day 1, reaching Stop 10 as quickly and safely as possible.
Stop 10. Santa Cruz Island Fault Overview (34.01760°N, 119.84029°W)
The SCI fault is the island’s largest surface fault and has the most evidence for recent earthquake activity. Its east-west trace across the island is well expressed in the topography (Fig. 2) and in the far west of the island, as seen from this stop, are very well-preserved drainage offsets. Northwest of Stop 10 and along the lower portion of the Northern Ridge, several left-lateral drainage deflections along the trace of the fault are clearly visible (Figs. 23 and 24). Offsets measure as much as 579 m (1910 ft) (Patterson, 1979; Pinter and Sorlien, 1991) and formed during the late Pleistocene and Holocene. (See topical note section on “Late Quaternary History of the Santa Cruz Island Fault.”) This is also a good stop to closely examine outcrops of the Willows Plutonic Complex, which is thought to be part of a remnant island-arc (Sorensen, 1985) and part of a larger collection of exotic terranes emplaced along the California Borderland. The complex consists mainly of hornblende gabbro and hornblende diorite with minor amounts of leucotonalite, gabbro, and ultramafics. There are many dikes that cross-cut the unit and some areas are intensely altered. (See topical note on “Exotic Terranes.”)
To drive to Stop 11, continue west to the floor of the valley while crossing the contact between lower Blanca Formation and Willows Plutonic Complex. Take the left fork near the airstrip and climb the ridge passing over Miocene Blanca and Rincon Formations, and Eocene Jolla Viejo Formation within the faulted north limb of a small anticline. The Jolla Viejo is a medium-grained, well-sorted arkosic sandstone that may have been deposited in a deep-marine channel complex (Boles, 2015). Within the coarse-grained beds of the Jolla Viejo Formation are meta-rhyolite clasts that resemble clasts from the Poway Conglomerate of the San Diego area (Yeats et al., 1974; Abbott and Smith, 1978, 1989). Trace-element analysis indicates the meta-rhyolite clasts in the Poway Conglomerate were sourced from Upper Jurassic bedrock in Sonora, Mexico (Abbott and Smith, 1989).
Along the ridge top, the road crosses southwest-dipping Vaqueros and Rincon Formations. The weakly resistant Rincon siltstone and shale are exposed in small gullies as the road descends into Canada de Sauces and crosses a syncline that bounds the northwest limb of the major Christi anticline. The road climbs out of Canada de Sauces and onto a ridge above the folded lower to middle Eocene Canada Formation along the complex northeastern limb of the Christi anticline. At the ridge crest is a good panorama to the south and southeast of the Canada Posa drainage and the west side of the Sierra Blanca that consists almost entirely of Blanca Formation. Our route descends into Poza Canyon and passes the degraded pad for a Richfield Oil Company (ROCO) exploratory well, the Santa Cruz Island #2 drilled in 1955. The well spudded in Paleocene Pozo Formation at the surface culmination of the Christi anticline, and encountered upper Cretaceous sandstone and conglomerate before reaching its total depth of 689 m (2260 ft).
Mapping of Dibblee (2001a) and Weaver and Nolf (1969) show the Christi anticline is a complex structure with two ages of folding, uplift, and erosion: (1) post–Cozy Dell Formation and pre–Vaqueros Formation deformation that occurred during the Oligocene and is recorded at the anticline as an angular unconformity; and (2) post–Blanca Formation (post–early Miocene) deformation. The pre-Vaqueros folding is of interest because it belongs to a set of regional mid-Tertiary unconformities that predate Miocene rotation and subsequent Transverse Range convergence. In the western Transverse Ranges, Dibblee (1950, 1976) referred to the folding and uplift as the Ynezian orogeny, and it is unclear from field relationships on SCI whether the deformation is due to crustal shortening or extension. (See topical note on “Mid-Tertiary Deformation.”)
Stop 11. Parking at Canada de Poza to Access the Wash
From the parking area (33.985364°N, 119.855114°W), we will make a 2–3-mile roundtrip hike down the wash to the Near Point section. This will be the lunch stop, too.
Stop 11B. Near Point Section (33.980251°N, 119.869265°W)
The Near Point section starts just west of the mouth of Posa Canyon and extends westward to Near Point (Figs. 25–28). The section is southwest dipping, along the southwest limb of the Christi anticline, and displays an early to middle Miocene sequence of Vaqueros, Rincon, San Onofre Breccia, and Beechers Bay Formations (Fig. 25). Along the east side of Near Point is an exceptionally well-exposed section of southwest-dipping San Onofre Breccia and overlying Beechers Bay Formation exposed in the sea cliff and modern wave-cut platform (Figs. 26–28). Closer to the mouth of Posa Canyon, the underlying Rincon and Vaqueros Formations occur but are less well exposed. The Beechers Bay Formation at Near Point is a distal facies to the lower or middle Blanca Formation that is exposed further east at Laguna Canyon (Howell et al., 1976; McLean et al., 1976; Boles, 2015). Bedding characteristics of the Beechers Bay Formation are similar to the thin-bedded strata of the Monterey Formation, and Weaver and Doerner (1969) assigned these rocks to the Monterey Formation. However, at Near Point, the Beechers Bay is a thin-bedded, very fine-grained volcaniclastic sandstone (Figs. 26 and 27) and compositionally much different from the siliceous, biogenic beds of the Monterey Formation. The San Onofre Breccia at Near Point consists of bedded conglomerates and sandstones with clasts of blue schist (glaucophane schist), dacite porphyry, quartzite, garnet amphibolite, and fossiliferous sandstone (Figs. 28A, 28B). (See topical notes on “Early–Middle Miocene Strata at Near Point–Upper Posa Canyon” and “The Monterey Formation.”) After viewing the Near Point section and lunch, return to the vehicle. If time permits, we will examine the San Onofre Breccia and Blanca Formation outcrops in upper Posa Canyon that are northeast of the parking spot. Depart Stop 11 and retrace road back to the ridge just north of Canada de los Sauces. Make a left onto a smaller dirt road and head west toward the coast, turn right, and head northwest toward Christi Beach (Stop 12). At the overlook above Christi Beach, the low-lying, very flat-topped, westernmost part of the island at Fraser Point can be seen. Fraser Point is capped by a late Quaternary marine terrace. (See topical note section on “Late Quaternary Marine Terraces.”)
Stop 12. Christi Beach Hike. Park at beach (34.022901°N, 119.876547°W)
We will make a 1–2-mile roundtrip hike southward along the beach to Christi Point. In the bluff just south of the parking spot is a syncline with Rincon Formation in its trough. South of the syncline axis is a northeast-dipping section of progressively older Rincon, Vaqueros, Cozy Dell, and Jolla Viejo Formations that are along the northeast limb of the Christi anticline (Fig. 29). Here the Vaqueros and Cozy Dell are separated by a low-relief angular unconformity.
Hike south along the beach approximately 1 mile to Christi Point, which is usually accessible. Then, examine the following stratigraphic section on the return trip north to the parking location.
Stop 12B. Upper Jolla Viejo, Cozy Dell, Vaqueros, and Rincon Formations Outcrop (34.016735°N, 119.879936°W)
At this stop, we will examine the lithology of the Jolla Viejo Formation. The Eocene Jolla Viejo Formation consists primarily of massive- to thick-bedded sandstone, interbedded with shale and shale- or volcanic-clast conglomerate (Figs. 30A–30B). Rip-up clasts, burrows, shale conglomerate, some cross-bedding, and dish structures suggest deposition within a submarine canyon or channel. The Eocene Cozy Dell Formation is a thin-bedded, fissile bluish- to greenish-gray, micaceous, foraminifera-rich siltstone/shale with thin limestone and sandstone beds. The foraminiferal assemblage indicates deep-water deposition. The shallow-marine lower Miocene Vaqueros Formation disconformably overlies the Cozy Dell, and consists of fossiliferous volcaniclastic sandstone and conglomerate (Figs. 30C–30D) containing principally volcanic and plutonic clasts with minor quartz-chlorite schist, but lacking the distinctive blueschist clasts of the younger, San Onofre Breccia. The Vaqueros Formation is conformably and gradationally overlain by the Miocene Rincon Formation, a gray-brown calcareous, foraminfera-bearing mudstone. Regionally, the Cozy Dell and Jolla Viejo Formations represent the latter stages of forearc basin development along Farallon–North American convergent margin during the Eocene. Transition from the extensive forearc setting to more localized transform basins is represented at SCI by the mid-Tertiary unconformity and overlying Vaqueros Formation.
Return to parking area and, time permitting, we will make a 1–2-mile hike north along the beach to Stop 12C.
Stop 12C. Exposure of Santa Cruz Island Fault (34.030946°N, 119.874981°W)
At Stop 12C is an ~25-m-high sea-cliff exposure of the Santa Cruz Island fault (Fig. 31) that juxtaposes Santa Cruz Island volcanics against nonmarine alluvial deposits of latest Pleistocene age (Pinter and Sorlien, 1991; Pinter et al., 1998b). This exposure indicates at least 25 m of vertical separation since the late Pleistocene (base of the nonmarine deposit is not exposed). The whitish fault zone dips steeply to the north, is several meters wide, and contains sheared and ground local bedrock that includes slivers of the Blanca Formation and Willows Plutonic Complex. Just above Stop 12C, Pinter et al. (1998b) mapped a series of trench walls dug into the nonmarine deposits that are cut by the SCI, in order to determine its latest Pleistocene and Holocene history. (See topical note section on “Late Quaternary History of the Santa Cruz Island Fault.”)
Hike back to parking spot and drive back to UC Field Station.
Day 4. Monday, 18 May 2020, 1.0 ft low tide @ 1:52 p.m. PDT
Prepare for the field and have breakfast. Today is the last day of the field trip, and we must clean up the portion of the UC Field Station we used, pack our gear into the vehicles, and depart the station at 9:30 a.m. for Stop 13.
Return to Canada del Puerto Road via the Canada del Medio Road connector. As we did on Day 2 of the trip, we will drive to Prisoners Harbor along the Canada del Puerto, make a right on the Navy Road, and continue along the road ~10 km (~6 miles) to Stop 13, which is a small turn around and parking spot at the end of Navy Road.
Stop 13. Parking at end of Navy Road (34.012920°N, 119.595350°W)
From here we will make a 2–3 mile roundtrip hike to the western flank of the Montañon Ridge, aka El Montañon highlands, to examine the basal Monterey contact with SCI Volcanics. (See the topical note section on the “Monterey Formation of Santa Barbara Channel” and the section on “Prehistoric Chert Use on East Santa Cruz Island” and its discussion of the “Santa Cruz Island Blonde chert.”) At the parking spot and along the old road to Montañon Ridge, the landscape is highly sculptured by rotated landslide blocks formed in the Monterey Formation (Fig. 32). Note the small areas of ponded drainage that form upslope of the rotated blocks.
Stop 13B. Tm/Tscv Contact (34.020870°N, 119.586620°W)
Along the west flank of Montañon Ridge, it is probable that the contact between Monterey Formation and the underlying Santa Cruz Volcanics is a low-angle detachment fault. Several observations support the fault interpretation: (1) Weaver and Nolf’s (1969) mapping of the Prisoners Harbor Member of the SCI Volcanics shows that it thins out below the base of the Monterey Formation as it heads southward across the isthmus (Fig. 33). Similarly, Dibblee (2001b) shows a marker bed in the Devils Peak Member intersecting the base of the Monterey Formation. (2) The lowest portions of Monterey Formation have a fractured and sheared appearance that increases toward the formation’s basal contact (Fig. 34). (3) Just above the contact, a thick chert breccia bed is present in the Monterey Formation (Fig. 35) and may be a fault breccia. Many of the basal chert beds and fragments show complete silica-replacement of originally carbonate sediments, including megascopic gastropods. (4) Such localized silicification also suggests abundant advective fluid flow that could take place along a detachment fault, as documented elsewhere at the base of the Monterey Formation (Wirtz, 2017; Davis, 2018).
Return to the vehicle and drive to Prisoners Harbor for return boat to Ventura Harbor and end of field trip.
For at least five decades, geologic data and observations from Santa Cruz Island (SCI) have been integrated with developing plate tectonic concepts that have played an important role in understanding convergent and transform plate boundaries processes, while explaining the perplexing distribution of rocks and structures in southern California. Geologic studies at SCI have also played an important role in earthquake hazards studies and more recently in studies on late Quaternary sea-level and climatic changes. The great amount and breadth of geoscience work done on the island make it very challenging for those new to the geology of the island and those trying to assimilate past work. These “Topical Notes” are an attempt to address these challenges by summarizing some of the key research on the island, and placing the field-trip stops into a broader context. It is beyond the scope of this guidebook to discuss the enormous amount of data, interpretations, and publications in any detail; so what follows are generalities from the literature and in places, the field-trip leader’s comments and interpretations.
Santa Cruz Island lies at the southwestern edge of the east-west–trending Transverse Ranges province (Fig. 36A) and west of the San Andreas fault plate boundary (Fig. 36B). Dibblee (1982a) divided the Transverse Ranges into three segments based on rock type and structural expression (western, central, and eastern) (Fig. 36A). Although lithologically distinct, the structure and geomorphology of the segments are the result of late Pliocene and Quaternary north-south shortening and most likely share a common lithospheric cause (Fig. 36C). Nearly 100 years ago, geologists recognized that the anomalous east-west–trending Transverse Ranges are both geomorphically and geologically unique compared to the northwest-trending provinces to the north and south, and required specific geologic explanation(s) (Hill, 1928; Reed, 1933; Reed and Hollister, 1936). The upper crust of the Transverse Ranges province is characterized by east-west–trending convergent faults and folds (Dibblee, 1982a; Morton and Yerkes, 1987), and compressive earthquakes such as the 1971 Sylmar, 1987 Whittier Narrows, and the 1994 Northridge (Davis and Namson, 1994) that result from, more or less, north-south–oriented principal stress. Initiation of the Transverse Ranges is not precisely defined and probably time-transgressive across the province: 4.0–2.2 Ma where the ranges meet the Los Angeles basin (Davis et al., 1989); 3.0–2.0 Ma where the northern edges of the western Transverse Ranges meet the San Joaquin basin (Davis, 1986); and for initiation of convergent structures within the Ventura basin (Yeats, 1983; Yeats et al., 1988; Namson and Davis, 1988b; Yeats and Rockwell, 1991).
Deeper geophysical studies show the Transverse Range province is underlain by a vertical slab of unusually high mantle velocities (Hadley and Kanamori, 1977; Raikes, 1980; Humphreys et al., 1984; Humphreys and Clayton, 1990; Zhao et al., 1996; Kohler, 1999). Under the central Transverse Ranges, a sizable crustal root has been modeled and quantified above the vertical mantle slab (Kohler and Davis, 1997; Kohler, 1999; Godfrey et al., 2002) (Fig. 36D). Sinking of the dense slab and downward pull of the overlying lower crust creates north-south shortening is expressed as convergent earthquakes and aseismic deformation in the upper crust above the brittle-ductile transition (Bird and Rosenstock, 1984; Namson and Davis, 1988b; Humphreys and Hager, 1990; Jackson and Molnar, 1990; Kohler, 1999). Two studies provide estimates of the amount of shortening across the Transverse Ranges, and the results are strikingly consistent despite the different approaches. Namson and Davis (1988b) constructed a retrodeformable cross section based on surface geology and oil and gas well data across the western Transverse Ranges, which calculated 53 km of north-south shortening in the upper crust across the entire western Transverse Ranges and San Andreas fault (Fig. 36C). Godfrey et al. (2002) modeled P-wave velocities from the 1994 Los Angeles Region Seismic Experiment that found an 8-km-thick crustal root beneath the surface trace of the San Andreas fault in the central Transverse Range, and allowed a mass balance calculation showing ~36 km of north-south shortening (Fig. 36D). The western Transverse Ranges shortening extends north of the San Andreas fault into the San Emigdio Mountains (Davis, 1986) accounting for an extra 19 km of shortening, while the San Andreas fault lies along the northern edge of the central Transverse Ranges. Removing the San Emigdio Mountains shortening gives 34 km of shortening in the western Transverse Ranges, compared to 36 km of shortening in the central Transverse Ranges. Vertical density anomalies can induce strong regional compressive stresses in the continental lithosphere without assistance of faraway forces (Fleitout and Froidevaux, 1982), such as rotating crustal blocks and restraining bends along large strike-slip faults (Fig. 36E).
The basement of the western Transverse Range segment consists of continental-crust plutonic and metamorphic rock in the east, and oceanic-crust ophiolite and accretionary wedge (Franciscan) rock in the west. Both basements are unconformably overlain by very thick sequences of Upper Cretaceous and Cenozoic sedimentary and volcanic rocks (Dibblee, 1982a). The Continental Borderland Province is characterized by northwest-trending highs and lows and active strike-slip faults (Marsaglia et al., 2019). Since both provinces are seismically active, their structural styles impart a strong influence on the landscape and offshore bathymetry that is easily recognized. Anacapa, Santa Cruz, Santa Rosa, and San Miguel Islands (the Northern Channel Islands) form an east-west–trending chain of topographic culminations along the southern edge of the Santa Barbara Channel (Reed and Hollister, 1936; Weaver et al., 1969). The chain is the westward extent of an antiformal uplift that extends westward from the Puente Hills, under the Hollywood and Elysian Park Hills (just north of metropolitan Los Angeles), the Santa Monica Mountains, and into the offshore (inset map, Figs. 1 and 3). The antiformal uplift has been interpreted to be the result of hanging-wall folding above a north-dipping, crustal-scale ramp in the ~200-km-long Elysian Park thrust system that is seismically active (Davis et al., 1989; Davis and Namson, 1994; Namson and Davis, 1991, 1996; Seeber and Sorlien, 2000; Sorlien et al., 2001). Shaw and Suppe (1994) and Pinter et al. (2003) referred to the system as the Northern Channel Islands thrust and the anticlinorium as the Northern Channel Islands anticline.
Several studies relate the development of the western Transverse Ranges to its continual clockwise rotation as a coherent block from early Miocene to present (Luyendyk et al., 1980; Jackson and Molnar, 1990; Luyendyk, 1991; Atwater, 1998). As described by Atwater (1998), at ca. 5 Ma, the Pacific plate captured Baja California and transported it northwestward where it rammed into southern California. The plate boundary shifted eastward in southern California, producing a left-step in the San Andreas fault and a restraining bend to right-lateral displacement. This model implies that north-south convergence across the bend caused uplift of the eastern part of the western Transverse Ranges segment, and the central and eastern segments of the Transverse Ranges, while continued clockwise rotation created convergence in the western Transverse Ranges. Rotation is needed in this model because the western Transverse Ranges extend to the west and beyond the San Andreas fault restraining bend, which is a lingering issue for relating western Transverse Range shortening to the restraining bend in San Andreas fault (Davis and Namson, 2017). Onderdonk (2005) studied the northwestern edge of the western Transverse Ranges, and concluded that clockwise rotation is recorded in the map-scale structures that show a “fan-like” closing process with an increase in north-south shortening to the west. However, retrodeformable cross sections constructed across the entire onshore portion of the western Transverse Ranges do not support a westward increase in shortening during the last 2.0–4.0 m.y. and do not show evidence for shortening in the Miocene (Namson and Davis, 1991, 1996). Levy et al. (2019) show a decrease in north-south shortening during the late Pliocene and Quaternary, from ~21 km near Ventura to 7 km near Point Conception using cross-section forward modeling. Despite this, relating late Pliocene and Quaternary shortening as recorded in convergent folds, faults, and earthquakes to clockwise rotation in the western Transverse Range, including SCI, is widely accepted.
Santa Cruz Island lies within the wide and complex plate boundary between the North America and Pacific plates (Figs. 36B). Snippets of the plate boundary history, from Jurassic to present, are recorded in the rock units and structures on the island. It is generally accepted that from Jurassic to late Eocene, an oceanic plate(s) was subducted eastward under the western edge of North America (Hamilton, 1969; Dickinson, 1981; Atwater, 1998). The dominant tectonic model since the late 1960s has the eastward-moving Farallon plate transporting various exotic fragments of continental and oceanic crusts and island-arc remnants, and accreting them to the western edge of North America (Hamilton, 1969). Sigloch and Mihalynuk (2013) offered a refinement that all proto-Pacific plates were consumed at stationary intra-oceanic subduction zones, and are now preserved as massive vertical slabs in the lower mantle that are revealed in deep geophysical studies and correlated to the shallow-level tectonic record. Their model proposes that subsequent crustal accretion occurred when western North America overrode the archipelago of abandoned subduction zones and island arcs, causing major episodes of Cordilleran mountain building. In these models, the SCI schist and Willows Plutonic Complex formed far from the North American plate, and were subsequently accreted to the North American plate before deposition of the Upper Cretaceous and lower Tertiary clastic marine rocks that are present on the southwest part of the island (Figs. 5 and 37).
Throughout coastal California from Point Arena south to the Santa Ana Mountains, pre-Oligocene rock units are overlain unconformably by middle–late Tertiary sequences (Dibblee, 1973, 1976, 1982b, 1982c). Much of the middle Tertiary record is missing from the San Diego area, but mid-Tertiary unconformities resume in northern Baja California. This break in the geologic record is shown by a set of time-transgressive unconformities that resulted from deformation, uplift, and erosion that initiated in Oligocene time, possibly late Eocene, and continued into the middle Miocene. In the southern and central Coast Ranges and western Transverse Ranges, middle Tertiary sequences commonly rest with significant angular discordance on an eroded surface of earlier deformed Upper Jurassic to Eocene strata and pre-Cenozoic plutonic and metamorphic rocks. Dibblee (1950, 1976) referred to the Oligocene deformation as the Ynezian orogeny, and he believed the scattered coarse-grained continental deposits of the Sespe, Simmler, and Berry Formations are the syn-orogenic deposits from the orogeny (Dibblee, 1973, 1982c). During the Oligocene, the East Pacific Rise was near the California portion of the western North American convergent margin, and resistance to the subduction of young, buoyant oceanic crust possibly explains the Oligocene deformation (Atwater, 1998), at least north of the present United States–Mexico border. Younger expressions of the mid-Tertiary unconformities in the western Transverse Ranges have been attributed to the beginning of rotation, at ca. 18 Ma (Onderdonk, 2005). However, the mid-Tertiary unconformities are much more regional than the rotations, and many of the unconformities have Oligocene and earliest Miocene strata above the unconformity (Dibblee, 1976) that are too old to be related to the start of Miocene rotations. Although evidence for Ynezian deformation is widespread, the structural style of deformation has received little attention and is open to study. Plate motion vectors indicate the Oligocene in central and southern California were a time of transtensional deformation (Atwater, 1998), but the Oligocene deformation recorded in map-scale structures in the western Transverse Ranges and San Emigdio Mountains is consistent with crustal convergence (Namson, 1987; Davis, 1986; Namson and Davis, 1988b).
Santa Cruz Island and the Progression of Plate Margin Ideas
Santa Cruz Island’s location along the southwestern edge of the western Transverse Ranges, and adjacent to the northern edge of the offshore Continental Borderland Province, provides a key field location to study the geology of southern California and plate margin processes (Fig. 37). In the 1960s and 1970s, tectonic kinematic models were presented to explain the dismemberment of Mesozoic to early Tertiary lithotectonic belts in southern California that are more coherent to the north and south, plus the widely separated outcrops of an Eocene depositional system containing the unique “Poway” clast-bearing conglomerates, and the scattered occurrence of the early Miocene San Onofre Breccia (Yeats, 1968, 1976; Jones et al., 1976). Key outcrops of the lithotectonic belts, the “Poway” clast-bearing conglomerate and San Onofre Breccia will be visited on the field trip. In the late 1970s, 1980s, and 1990s, volcanic and sedimentary rocks on SCI played an important role in the discovery that the island and the western Transverse Ranges had rotated ~90° clockwise since the early Miocene (Kamerling and Luyendyk, 1979; Luyendyk et al., 1980; Hornafius, 1985). Clockwise rotation of the Miocene rocks of the western Transverse Ranges (Fig. 38) is now broadly accepted in the geologic community (Fritsche et al., 2001; Hillhouse, 2010), and resulted in several models having the entire western Transverse Ranges rotate as a coherent block with large strike-slip faults to the south of SCI (Kamerling and Luyendyk, 1979, 1985; Crouch, 1981; Hornafius, 1985; Hornafius et al., 1986; Luyendyk and Hornafius, 1987; Luyendyk et al., 1980, 1985; Jackson and Molnar, 1990; Meigs and Oskin, 2002). These rotation/strike-slip models account for the disrupted lithotectonic belts, the separation and distribution of “Poway” clasts and San Onofre Breccia, plus the separation of very similar Paleocene stratigraphic sections in Santa Monica and Santa Ana Mountains of mainland southern California, Paleocene paleocurrent directions, and the cause of the extensive early to middle Miocene volcanic activity in a postsubduction setting (Fritsche et al., 2001). However, as previously discussed in the topical note on “Transverse Ranges,” it is debatable if clockwise rotations played a role in development of the present Transverse Ranges that occurred during the last 2–4 m.y.
Howell et al. (1976) and Howell and Vedder (1981) proposed that large right-lateral displacements of far-formed tectonostratigraphic terranes along borderland faults could account for the disrupted lithotectonic patterns and incorporate local rotation (Fig. 37). SCI is a key location for Howell and Vedder’s (1981) terranes, and the SCI fault is the boundary of their Terranes II and III (the field trip will visit some of the rocks of Terrane II: SCI Schist, Alamos Tonalite, Willows Plutonic Complex, and Tertiary marine strata; and some of the rocks of Terrane III: Miocene volcanic rocks may have been encountered in an exploratory well on the northeast part of the island). As pointed out by Bohannon and Geist (1998), the rotation and strike-slip models account for much of the previously mentioned complexities, but the models do not account for the abbreviated rock column of the southwestern Los Angeles basin and inner borderland, where Miocene strata and volcanics rest directly on Catalina Schist (Fig. 37). The missing part can be explained by large-magnitude extension, accompanied by significant uplift and exposure of the Catalina Schist into the space opened behind the rotating western Transverse Ranges (Kamerling and Luyendyk, 1979; Legg, 1991; Sedlock and Hamilton, 1991; Crouch and Suppe, 1993; ten Brink, 2000). Bohannon and Geist (1998) strengthened the rotation and rifting model by adding a considerable amount of offshore geologic and geophysical data that tie the geometry and kinematics of the model to many of the principal geologic structures of the borderland. Nicholson et al. (1994) present an interpretation that rotation is due to the Pacific plate capturing the partially subducted Monterey micro-plate ca. 20 Ma, coupled with a change in the vector of Pacific plate motion. Accretion of the Monterey plate to the Pacific plate created a right-lateral transform boundary that initiated as a system of low-angle faults that produced extension in the overlying continental margin. Northward drift of the combined Pacific-Monterey plates beneath the margin produced right shear that resulted in clockwise rotation (Fig. 38). Atwater (1998) and Atwater and Stock (1998) provide comprehensive but condensed accounting of this enormous amount of research.
While large clockwise rotation of Miocene rock is well documented, the rifting and rotation model(s) described above raises questions. As Dibblee (1982a) noted, rotating the western Transverse Ranges 90°, or even more, as a coherent block presents space problems with adjacent blocks that are difficult, if not impossible, to account for in the geologic mapping. Dibblee may have been referring to the lack of map-scale structures, convergent or extensional, that would document Miocene deformation. Dickinson’s (1996) model of fault accommodation may offer a solution to Dibblee’s question, but needs to be evaluated at map-scale. Large rift systems of a similar scale to the Inner Borderland rift, such as the Red Sea, Death Valley, and Gulf of California, record a distinct set of extensional structures along their margins. For the most part, map-scale extensional structures seem absent from the onshore portions of the western Transverse Ranges (including SCI) or have received little recognition or study. Schwartz (2018) suggests that Miocene rotations occurred as separate microplates, rather than a coherent block, between 12 and 8 Ma and well before development of the western Transverse Ranges fold and thrust belt.
The central anticline has a N80W trend (measured along the SCI fault), a north limb with 10° to 30° northeast dips, and a steeper south limb with 15° to 60° dips (Dibblee, 2001a, 2001b). The anticline has no obvious hinge domain or rollover, and the SCI fault is the dividing line between the north- and south-dipping limbs along the entire onshore fault trace. The origin of the central anticline and even calling it an anticline are debatable. Occupation of the hinge area by the SCI fault begs the question: Did the structure form as a single fold, or did recent strike-slip movement on the SCI fault bring together two opposing anticlinal limbs, that grew initially along two separately formed anticlines? Those preferring the later origin may not consider the structure an anticline, if they consider the term to imply its deformational cause. The Glossary of Geology (Bates and Jackson, 1980) offers this description of an anticline: “a fold, generally convex upward, whose core contains stratigraphically older rocks.” It appears that “anticline” is a descriptive term. Plus, anticlines form in a variety of deformational settings: convergent, strike-slip, and extensional. The strike and dips, and stratigraphy recorded in the mapping of Weaver and Nolf (1969) and Dibblee (2001a, 2001b) indicate that the antiformal structure abreast the SCI fault (central anticline) meets the descriptive definition of an anticline. Furthermore, the central anticline is shown as an anticline in cross sections in at least two published papers (Fig. 3B) (Shaw and Suppe, 1994; Pinter et al., 2003).
The cross sections in Figure 3B show the central anticline is either the subaerial expression of the Northern Channel Islands anticlinorium (Pinter et al., 2003), or the central anticline is a separate smaller and more complex structure developed above the anticlinorium (Shaw and Suppe, 1994). Similarly, the Christi and northeast anticlines of SCI are developed above the anticlinorium. Both cross sections show the anticlinorium caused by a ramp (bend) in the surface of the north-dipping fault (Northern Channel Islands thrust, i.e., the western continuation of the Elysian Park thrust of Davis et al., 1989). Alternatively, Baker (1993) interpreted the south limb dip to be the result of Miocene extension that produced hanging-wall block rotation. In this interpretation, large north-dipping normal faults were active during the Miocene along the north and south edges of what was to become the south-dipping limb. The northern fault is the ancestral SCI fault, and the southern fault is located offshore along Kelez ridge, which is southwest of Santa Cruz and Santa Rosa Islands. Down-to-the-north movement on these faults caused hanging-wall block rotation and the south limb dip.
The Shaw and Suppe (1994) and Pinter et al. (2003) cross-section interpretations in Figure 3B show the central anticline and deeper thrust are convergent structures that resulted from late Pliocene and Quaternary north-south shortening associated with development of the western Transverse Ranges. Offshore 2D reflection seismic images and onshore geologic relationships to the east indicate the Northern Channel Islands anticlinorium developed during the late Pliocene and Quaternary (Davis et al., 1989; Pinter et al., 1998a, 2003; Seeber and Sorlien, 2000; Sorlien et al., 2001). Pinter et al. (2003) measured uplifted marine terraces and submerged shelf-edge deltas on and around Santa Cruz and Anacapa Islands, and show the foci of uplift at SCI to be under the north limb of the central anticline and broadly across the northeast anticline of SCI. From this work, Pinter et al. (2003) calculate that post–400 ka uplift of the anticlinorium is 0.7 mm/yr with the limbs subsiding up to 0.8 mm/yr.
Late Quaternary folding and uplift on the Northern Channel Islands anticlinorium, and coeval strike-slip movement on the SCI fault may be a unique example of strain partitioning on an individual map-scale structure (this is implied in the cross sections of Shaw and Suppe, 1994; Pinter et al., 2003; Fig. 3B). For simplicity, strain partitioning is defined here as the separation of deformation into pure and simple shear components that can occur at various scales. A good local example is recorded near the San Andreas fault in central and southern California (Davis and Namson, 2017). There, the transpressive stress along the plate boundary has resulted in partitioned structures: (1) convergent pure shear expressed as small fold and thrust belts parallel, or sub-parallel, to the trace of the San Andreas fault; and (2) simple shear expressed as nearly pure strike-slip offset features along the San Andreas fault with no sign of long-term vertical offset recorded in the surface or subsurface geology. Zoback et al. (1987) and Mount and Suppe (1987) showed that present-day in situ tectonic stress indicators along this segment of the San Andreas fault are at a high angle to the trace of the fault, and deformation is not controlled by distributed shear associated with a high-drag San Andreas fault. Both studies conclude that transpressive tectonics along the plate boundary can be better described as decoupled transcurrent and compressive deformation, operating simultaneously and largely independently. Fault-normal crustal compression results from the extremely low shear strength of the San Andreas fault, and the slightly convergent relative motion between the Pacific and North American plates. Clearly, the structural origin, geometry, offshore extent, and age of the central anticline and Northern Channel Islands anticlinorium, and their relationship to the SCI fault are subjects that need further research.
Santa Cruz Island Volcanics
The SCI Volcanics are part of a more regional episode of volcanism that spread across the Channel Islands and coastal California at 19–13 Ma (e.g., Weigand, 1993, 1997; Weigand and Savage, 2002). The SCI Volcanics range in composition from mafic to intermediate to silicic as the sequence progresses and is divided into four members: Griffith Canyon, Stanton Ranch, Devils Peak, and Prisoners Harbor (Fig. 4). Within each member, there are numerous interbedded flows, flow breccias, and volcaniclastics. The ages of each member overlap, probably due in part to multiple volcanic edifices. The lower members are dominated by more mafic compositions, while the middle and upper members are dominated by intermediate compositions. The base of the SCI volcanic sequence is inferred from the Richfield Oil Company #1 well, indicating a cumulative thickness of ~2500 m (~8202 ft) (Nolf and Nolf, 1969), and is presumably in contact with the San Onofre Breccia. The lowermost three members of the SCI Volcanics erupted subaerially, but the later Prisoners Harbor units erupted subaqueously (Nolf and Nolf, 1969). The SCI Volcanics are overlain by the Monterey Formation (Fig. 4). Much of the SCI Volcanics appear to have erupted locally, and therefore, the unit boundaries are complex, involving preexisting topography (Nolf and Nolf, 1969). The western half of the volcanics is cut by intrusives, which may be the source of the volcanics (Weigand, 1997).
The SCI Volcanics are dominantly calc-alkaline with no evidence of extensive crustal contamination. The REE (rare earth element) profiles show moderate light rare earth element enrichment with small Eu anomalies (Johnson and O’Neil, 1984; Weigand, 1987). Sr, O, and Nd data suggest a mantle source in a subduction environment (Weigand, 1987). Overall, the SCI Volcanics are more enriched in REE than the Conejo Volcanics in the Santa Monica Mountains on the mainland (Weigand, 1997), but early studies suggested they may be correlated (e.g., Nolf and Nolf, 1969).
Because of limited geochemical studies of the SCI Volcanics, the mechanism of magma generation remains elusive. Weigand (1997) provides a succinct summary of the leading theories. Some authors have attributed the volcanism to arc magmatism related to the subduction of the Farallon plate (e.g., Weigand, 1982; Crowe et al., 1976). Others have attributed the magmatism to the subduction of the Pacific-Farallon spreading center (Dixon and Farrar, 1980). Finally, there is mantle diapirism due to a slab-free window (Dickinson and Snyder, 1979; Wilson et al., 2005) or the possible presence of a weak hotspot (Day et al., 2019).
Early–Middle Miocene Strata at Near Point–Upper Posa Canyon
The Near Point section and nearby sections are described in detail in Boles (1997), Boles and Weigand (1997), and Boles (2015), and the following is a summary of these works, plus additional field observations by Richard Behl and Thomas Davis. The Near Point section starts just west of the mouth of Posa Canyon and extends westward to Near Point. The section is southwest dipping, along the southwest limb of the Christi anticline, and displays an early to middle Miocene sequence of Vaqueros, Rincon, San Onofre Breccia, and Beechers Bay Formations (Figs 5 and 39). The Vaqueros rests unconformably on Eocene Cozy Dell Formation (Dibblee, 2001a), but this key contact and the Vaqueros Formation are not well exposed here. Along the northeast limb of the Christi anticline, Dibblee (2001a) shows the basal Vaqueros unconformity as angular, indicating Oligocene deformation. The upper part of the Rincon Formation, San Onofre and Beechers Bay Formations are very well exposed in the sea cliff and wave-cut platform just east of Near Point, and the near perpendicular exposures between cliff and platform offer an almost 3D perspective of sedimentary bedding and structures (Figs. 25 and 26).
Where the Vaqueros Formation is well exposed locally at Canada Alegria, the formation is up to 240 m (787 ft) thick and composed of a lower conglomeratic sandstone and an upper lithic sandstone rich with volcanic debris. Large cross beds, rare clast imbrications, abundant marine fossils (pelecypods, gastropods, crustacea, and shark teeth), plus scallop and oyster beds suggest a shallow water environment of deposition. At Near Point, the early Miocene Rincon Formation conformably overlies the Vaqueros Formation (Fig. 39). There, the Rincon is 76 m (249 ft) thick, but it is only 20 m (66 ft) thick in upper Posa Canyon, where its upper part is eroded by the overlying San Onofre Breccia. The Rincon Formation consists of gray to brown calcareous mudstone. Overlying the Rincon Formation at Near Point is 99 m (325 ft) of early Miocene San Onofre Breccia that includes conglomerate, sandstone, and mudstone beds. There, the San Onofre is mostly medium-bedded conglomerate and sandstone that are in contrast to the thick conglomerate/breccia beds in upper Posa Canyon. Clasts within the San Onofre Breccia consist of dacite porphyry, quartzite, garnet amphibolite, glaucophane schist (blue schist) and fossiliferous sedimentary rocks. The San Onofre Breccia occurs throughout coastal southern California and northwestern Baja California, and its distinctive clast assemblage (an abundance of blue schist and garnet amphibolite schist clasts) is thought to be derived from the Catalina Schist terrane that is subaerially exposed on Catalina Island and the Palos Verdes Peninsula (Woodford, 1925; Stuart, 1979). Deposition of the Near Point outcrops of San Onofre Breccia was probably in an inner-shelf setting based on sparse assemblages of pelecypods. Overlying the San Onofre Breccia at Near Point is 177 m (581 ft) of Beechers Bay Formation that consists of conglomerate, sandstone, and often intraformationally deformed, thin-bedded siltstone. Compositionally it is volcaniclastic and geochemically and stratigraphically equivalent to the thick tuffs of the Blanca Formation further east in Laguna Canyon. Foraminiferal assemblages indicate deposition occurred in increasingly deeper waters, from outer shelf to upper slope (Bereskin and Edwards, 1969).
As noted by Boles (1997), the San Onofre Breccia in upper Posa Canyon (near our parking spot at Stop 11) is remarkably different from the section at Near Point despite being only 3.2 km (~2 mi) apart (Fig. 39). In the upper canyon, the breccia/conglomerate beds are thick and contain large clasts, some over a meter in size. Sandstone beds are rare and occur at the top of upward-fining cycles. In places, oyster fossils are attached to larger clasts in growth assemblages within sandstone beds, and probably indicate very shallow water conditions, possibly less than a few meters. Abundant, imbricated schist clasts show paleocurrent directions to the west-southwest. Based on grain size and paleocurrent indicators, the San Onofre Breccia at Near Point appears distal to the San Onofre Breccia in upper Posa Canyon. Boles (1997) interprets the southwest transition to be in the direction of sediment transport, and blue-schist dominated facies in upper Posa Canyon to be derived from a high-relief source, possibly a fault scarp. Also noted by Boles (1997) is the inconsistency of west- to southwest-directed sediment transport with subsequent clockwise rotation of the western Transverse Ranges. Removing the +90 degrees of postdepositional clockwise rotation places the Catalina Schist source area to the north and northwest, which does not easily fit accepted tectonic models. Rather, rifting of the western Transverse Ranges (and SCI) from the Orange County–San Diego area should have exposed the Catalina Schist from a broad easterly direction during the early Miocene, not from the west and northwest. Admittedly, sediment transport directions can vary locally, especially in tectonically active areas; plus map relationships along the northeast limb of the Christi anticline suggest structural growth during deposition of the lower Blanca Formation during the Miocene (Dibblee, 2001a), and fold growth could have influenced local depositional patterns.
The Monterey Formation of Santa Barbara Channel
The Miocene Monterey Formation is perhaps the most globally recognized and important formation in California. It is the primary source of oil in the state and an important petroleum reservoir (Isaacs and Rullkötter, 2001). It was deposited across the critical Miocene climatic transition when permanent ice sheets were established in Antarctica, and global oceanic circulation patterns changed to bring increased upwelling of nutrients to the surface of the Pacific Ocean (Ingle, 1981; Barron, 1986; Pisciotto and Garrison, 1981). Abundant siliceous, calcareous and carbonaceous sediments accumulated within the Monterey Formation and its equivalents around the Pacific Rim during the middle to late Miocene, and the Monterey became the focus of numerous studies and much of our understanding of the depositional and diagenetic processes of these kinds of fine-grained, hemipelagic sediments (see review in Behl, 1999). Many of the key studies took place in the Ventura–Santa Barbara basin (cf. Isaacs, 1981, 1985), the southern boundary of which is formed by the Santa Cruz Island–Anacapa extended fault system. Monterey strata crop out north of the SCI fault on the island and, along with the SCI Volcanics, make up the southernmost deposits of the Neogene sedimentary basin.
Although widely known as an organic-rich, highly siliceous deposit that recorded the accumulation of abundant diatomaceous sediment, stratigraphic heterogeneity in the Monterey Formation has been documented for more than 80 years. Some of this heterogeneity is due to primary compositional differences that form the basis of official or unofficial lithostratigraphic members, and some is due to expression of the diagenesis of siliceous sediments (Isaacs, 2001). The Monterey is divided into different members with different names in the various basins or parts of basins of California, but many of these schemes share a general compositional theme of being more calcareous in the lower portions and more organic rich in the middle parts (Fig. 40). Primary compositional variation is complicated and sometimes obscured by silica diagenesis, in which originally diatomaceous sediment—highly porous, pure diatomite to slightly diatomaceous mudstone containing opal-A—transforms to denser, harder chert, porcelanite, or siliceous shale containing two distinct diagenetic polymorphs of silica—opal-CT and quartz. The interrelationships between original sediment composition, diagenesis, resultant rock properties, and styles of tectonic deformation are critical to understanding the history of the Monterey and to exploiting it for its resources (Schwalbach et al., 2009; Behl and Gross, 2018). No work has been done on SCI since development of this modern integrated perspective.
Less appreciated than the vertical stratigraphic variation in the Monterey are the significant lateral variations, even within the fine-grained lithofacies. Hornafius (1991, 1994) demonstrated an approximately five-fold expansion of certain lithostratigraphic units in the Santa Barbara basin over just 15–30 km (~9.3–18.6 miles), which was accomplished by increased accumulation or focusing of originally diatomaceous sediment into local depocenters. Doubling of stratigraphic thickness of units in the Antelope and McDonald Shale members of the Monterey Formation over just a few kilometers in the Belridge oil field (San Joaquin basin, California) occurs with dramatic variation in clay, silica, organic carbon content, and downhole log character (Farrell, 2020). Different patterns of sediment accumulation may be due to a number of variables, including water depth and effects of current winnowing, downslope of along-axis gravity flow redistribution, proximity to sources of clastic sediments, and dissolution of soluble phases (Fig. 41). Schwalbach and Bohacs (1992) proposed a sequence-stratigraphic model for variation in Monterey sedimentation that reflected both bathymetric position and deposition within parts of a sea-level or climate-oceanographic cycle. With all these indications of lateral and vertical variability, there is little reason to expect that the Monterey Formation on SCI would closely compare with any particular stratigraphic section on the mainland or in the subsurface of the Santa Barbara Channel. Yet, there has been remarkably little work done on the Monterey Formation on the island.
The foundational work on the SCI Monterey Formation was by Rand (1933), whose Ph.D. dissertation covered the geology of the entire island! More than 80 years later, his work is still the most detailed lithostratigraphic treatment of the Monterey Formation, and his work is generally cited by subsequent authors. Rand described the basal contact with the SCI Volcanics in a number of locations and >478 m (1570 ft) of interbedded chert, limestone (dolostone), and foraminiferal shale that was separated into four stratigraphic intervals. The most complete succession extends from the basal contact at Montañon Ridge westward to the axis of the first northwest-southeast–trending syncline (that is exposed in Chinese Harbor). Yet, these descriptions were done before the modern understanding of the effect of silica diagenesis on lithology had been developed, which started with Bramlette (1946) and advanced in the 1980s and 2000s with many workers. Weaver and Meyer (1969) extended Rand’s work by identifying an enigmatic chert breccia at the basal contact at Montañon Ridge, and the discontinuous nature of some of the calcareous units near the base of the section. They also improved the benthic foraminiferal biostratigraphy to identify the age of the SCI Monterey as lower Relizian to upper Luisian stage—which could span from 17.1 to 13.5 Ma in the updated chronostratigraphy of Barron and Isaacs (2001). This is the same age as the lower part of the sections at El Capitan and Naples beaches on the Santa Barbara coast, and almost exactly the age range (terminating in a similar tectonic erosional unconformity) that occurs at Shell Beach (Föllmi et al., 2017).
The stratigraphy of the Monterey Formation is difficult to correlate over much of SCI due to poor exposure and the abundance of landslides that formed in the unstable, thin-bedded rocks. The best exposures are in the sea cliffs and the drainages west of Montañon Ridge, where we will make three trip stops.
With so little work done, there are many things to investigate about the Monterey Formation on SCI:
What is the diagenetic phase of the rocks and how does that vary, both stratigraphically and laterally, through the island?
What is the burial history of the island that can be determined by sedimentary geothermometers?
How does the depositional environment of the basal contact vary in space, and what does that say about the transition from SCI Volcanics to Monterey?
What is the significance of the chert breccia and extensive silicification at the basal contact at Montañon Ridge? Does it relate to a tectonic detachment seen elsewhere in California?
Are fault contacts between the SCI Volcanics and Monterey syn- or post-depositional?
Can burnt shale deposits be studied in 3D space on the island to better understand their genesis?
(See section on “Prehistoric Chert Use on East Santa Cruz Island.”)
Late Quaternary History of the Santa Cruz Island Fault
The late Quaternary displacement history of the SCI fault has been studied through landform analysis and trenching of the fault trace (Patterson, 1979; Pinter and Sorlien, 1991; Pinter et al., 1998a, 1998b). The following is summarized from Pinter et al. (1998a, 1998b): Mapping of a trench site located above the sea cliff at Christi Beach, and nearby terrace deposits indicate multiple large-magnitude earthquake events have occurred on the fault during the late Pleistocene and Holocene with the latest event ca. 5 ka. Displacement of adjacent marine terraces, dated by uranium series, and other landforms show an average left-lateral slip rate of 0.8 mm/yr and reverse slip of 0.1–0.2 mm/yr. Average late Quaternary recurrence interval is estimated to be at least 2.7 k.y. and probably between 4 and 5 k.y. with earthquake magnitudes of 7.2–7.5 Mw. These estimates are similar to other faults along the southern margin of the western Transverse Ranges, i.e., Malibu Coast, Santa Monica, Hollywood, and Raymond faults, suggesting a connected system of faults that is characterized by large infrequent earthquakes.
Late Quaternary Marine Terraces
Santa Cruz Island and several of the other Channel Islands of southern California preserve an exceptionally good record of major late Quaternary sea-level fluctuations as a result of major climatic changes. Understanding periods of higher than present sea level and their impacts are currently of great societal interest, and recognized as important subjects for future research by the latest report of the Intergovernmental Panel on Climate Change (Church et al., 2013; Masson-Delmotte et al., 2013). The last interglacial (LIG) is considered by many to be a relevant analog for the future, as the mass of polar ice was significantly less during the LIG and sea level was higher than today (Murray-Wallace and Woodroffe, 2014). In addition, the elevations and ages of late Quaternary terraces are used to estimate tectonic uplift rates that are important to seismic risk evaluation.
The present elevation of marine terraces along the Pacific Coast of North America that formed during the LIG (~120 ka) are the result of three coeval processes: eustatic sea-level changes, glacial isostatic adjustment (GIA), and vertical tectonic movements that are common to the Pacific–North American plate margin. Worldwide eustatic sea level during the LIG is considered to be between ~3 and ~10 m above present level with 6 ± 3 m commonly quoted (Muhs and Groves, 2018; see review in Murray-Wallace and Woodroffe, 2014). These values match well with LIG terraces at Isla Guadalupe off western Mexico, located far enough from the plate boundary to be tectonically stable. Any departure from the ~3 to ~10 m above present sea-level elevation for LIG terraces is a sign of uplift or subsidence. GIA results from large ice sheets covering large portions of North America as recently as 20 ka, with ice sheet loading and removal affecting the crust and lithosphere. Marine terraces have significant uplift at locations occupied and adjacent to former ice sheets, and uplift decreases with distance from paleoloads. Coastal California, while not under an ice sheet in the late Quaternary, occupies an “intermediate field” for GIA effects (Muhs and Groves, 2018). Modeled estimates of GIA for the LIG along the California coast are ~11 to ~13 m above present sea level (Creveling et al., 2015) to 12 to 13 m above present sea level (Simms et al., 2016). Understanding GIA is key to determining the impact of rising seas during the LIG and estimating tectonic uplift rates for seismic risk evaluations. How well the modeled rates match the field evidence is discussed below.
Along the high-energy Pacific coastline of North America, sea-level changes are recorded in the geologic record as marine terraces commonly preserved in a staircase-like coastal topography. Terraces are remnant wave-cut platforms that formed during interglacial periods of high sea-level stands with a thin covering of mostly non-marine sand and gravel. Maximum paleo-sea levels are most precisely recorded by shoreline angles that form where the ancient sea-cliff and wave-cut platform intersected. Detailed mapping and precise elevation measurement of the shoreline angle, in conjunction with radiocarbon, uranium-series, and amino-acid dating methods of marine fossils, have been key to building a detailed chronology of climatically driven eustatic sea-level changes and determining tectonic uplift rates.
Santa Cruz Island was uplifted during the Quaternary, as is much of coastal California, and where geomorphic preservation is favorable, late Quaternary marine terraces remain. Well-preserved marine terraces in westernmost SCI near West and Fraser Points, and between Kinton Point and Punta Arena were mapped and measured in detail by Pinter et al. (1998a, 2003) (Fig. 42). Pinter et al. identified three terrace levels: T1, T2, and T3 (from youngest to oldest). Uranium-series ages of fossil corals from T1 near Near Point and above Stop 11B (Fig. 25) are ~120 ka, i.e., the LIG (Pinter et al., 1998a). LIG is within the range of MIS 5.5, or 5e (Shackleton and Opdyke, 1973) that occurred ~130 ka to ~115 ka (MIS—marine isotope substage of the foraminiferal oxygen isotope record of deep-sea cores [Martinson et al., 1987]). Close to Pinter et al.’s (1998a) fossil coral sample location, amino-acid dating of the gastropod Chlorostoma plot within the ~120 ka or 120 ka to 100 ka aminozones (Muhs and Groves, 2018), consistent with the nearby U-series age.
At Fraser Point, the T1 surface (youngest marine terrace, Pinter et al., 1998b) is 6–8 m above present sea level, but its deposits lack corals, one of the few marine invertebrates that incorporate uranium from seawater into their skeletons during growth. Amino-acid dating of fossil mollusks is an alternative dating method that provides relative ages that can be tied over distance. Amino-acid dating of Chlorostoma from T1 deposits at Fraser Point places the deposit in the ~120 ka or 120 ka to 100 ka aminozones. At Fraser Point, the next older terrace is T2 and it is 20–40 m above present sea level, and T3, the oldest terrace, is ~120 m above present sea level. Both of these higher marine terraces are undated and older than T1, as the island has a history of continual uplift during the Quaternary.
The shoreline angle elevations at Fraser Point, integrated with the modeled GIA estimates of paleo–sea level during the LIG by Creveling et al. (2015) and Simms et al. (2016) require ~3 to ~6 m of subsidence since the LIG high-sea stand, and tectonic subsidence is not considered valid due to the tectonic setting and history (Muhs and Groves, 2018). However, one could make a very speculative case that young subsidence is occurring in westernmost SCI. The traces of the Santa Cruz and Santa Rosa Island faults form a releasing left step in a left-lateral fault system that has well-documented strike displacements during the late Quaternary (Dibblee, 1982b; Pinter et al., 1998b). However, it is more likely the modeled GIA estimates of Creveling et al. (2015) and Simms et al. (2016) since the LIG are too high, given the LIG terrace elevations at Fraser Point, on nearby southern California islands, and at Isla Guadalupe (Muhs and Groves, 2018).
Quaternary tectonic uplift rates are key to seismic risk evaluation in regions undergoing active convergence, based on the assumption that the higher the rate, the more frequent the earthquakes. Uplift rates are an especially critical measurement in convergent belts with blind thrusts such as in the Transverse Ranges and Coast Ranges of California, where surface rupture observations are absent (Namson and Davis, 1988a, 1988b; Davis et al., 1989). At SCI and the other Northern Channel Islands published uplifts vary greatly. Sorlien (1994) postulated an uplift rate of 0.5–1.0 m/k.y. for the Northern Channel Islands, based on terraces at Santa Rosa Island. Pinter et al. (2003) estimated the anticlinorium forming Santa Cruz and Anacapa Islands has an uplift rate of 0.7 m/k.y. at its crest with the submerged limbs subsiding at a rate of 0.8 m/k.y., for the last 400 k.y. In comparison, uplift rates for the Northern Channel Islands that incorporate the GIA effects are lower (e.g., Muhs et al., 2012). San Miguel and Santa Rosa Islands range from 0.12 to 0.15 m/k.y. since the LIG (Muhs et al., 2014). The elevation and age of the LIG terrace (T1) on western SCI (Pinter et al., 1998a, 1998b, 2003) indicate a maximum rate of ~0.09 m/k.y., and possibly no uplift at all in the northwesternmost part of the island (Fraser Point), according to Muhs and Groves (2018). Furthermore, Muhs and Groves’ (2018) uplift rates are an order of magnitude less than those of Chaytor et al. (2008), whom calculated a 1.5 m/k.y. rate during the last ~23 k.y. by mapping and dating a submerged terrace just south of the Northern Channel Islands, which developed during the last glacial maximum. Muhs and Groves (2018) explain that the high uplift rate of Chaytor et al. (2008) is an artifact that did not account sufficiently for the significant GIA effect along the California Coast since the LIG (Muhs et al., 2012).
THE FLORA AND FAUNA OF SANTA CRUZ ISLAND
Flora of Santa Cruz Island
Isolated off Santa Barbara, Santa Cruz Island is home to 60 endemic species found no place else in the world. Some of these species include the Santa Cruz Island bushmallow (Malacothamnus fasciculatus var. nesioticus) and the island malacothrix (Malacothrix squalida). Below is an introduction to some of the flora that can be found on the island, organized by the principal geologic unit in which they can be found.
Santa Cruz Island Schist. The SCI Schist hosts a wide variety of plant communities including “chaparral, coastal-sage scrub, grasslands, island woodland, pine forest, riparian woodland, riparian scrub and oak woodland” (Junak et al., 1995, p. 10). Some of the indicator species of the chaparral are manzanita plants (such as Arctostaphylos insularis and A. tomentosa subsp. insulicola).
In the woodlands, look for the Santa Cruz Island ironwood trees (Lyonothamnus floribundus subsp. aspleniifolius). These are a remnant species that were once found on mainland California and now only exist naturally on the Channel Islands. The ironwood trees tend to grow at a certain altitude to take advantage of the prevalent fog and to avoid the temperature extremes of the inland valleys.
Santa Cruz Island Volcanics. “Santa Cruz Island Volcanics support a variety of tree dominated plant communities” (Junak et al., 1995, p. 11). The canyons and north-facing slopes provide excellent habitat for oak trees (such as Quercus agrifolia and Quercus tomentella), willow (Salix spp.) and maple (Acer macrophyllum).
Two beautiful native plant species can also be found on the SCI Volcanics; the Santa Cruz Island buckwheat (Eriogonum arborescens) and Northern Island hazardia (island bristleweed) (Hazardia detonosa) (Figs. 43 and 44). The range of these plants is restricted to Santa Cruz Island and some of the other Northern Channel Islands.
Also found in the volcanics are prickly pear cactus. There are several species (including Opuntia littoralis, O. oricola, and a hybrid) found on Santa Cruz Island, and one way they can be distinguished are by their leaf pad shapes. The fruits of the Opuntia are edible, though laden with spikes. The Opuntia also hosts the cochineal insect, an insect that can be used as a source of dye for clothing.
As you walk through Canada del Puerto, keep an eye out for the low-growing Santa Cruz Island silver lotus (Acmispon argophyllus var. niveus) (Fig. 45). This rare plant is endemic to SCI. The SCI Volcanics are also home to other endemic species, including the Santa Cruz Island live-forever (Dudleya nesiotica)—a rare succulent whose range is confined to a small portion of the island.
Monterey Formation. “Monterey Shale primarily supports grassland, with scattered stands of chaparral dominated by scrub oaks, coastal-sage scrub, and an extensive Bishop pine forest” (Junak et al., 1995, p. 12). Monterey shale is “relatively impermeable to water” and the area is more prone to slides, which in turn affects the “soil formation and may inhibit the establishment of some plant communities” (Junak et al., 1995, p, 12). The rare white-haired manzanita (Arctostaphylos viridissima) can be found on the Monterey Shale.
One of the indicator species of coastal-sage scrub is coastal sagebrush, which is also known as California sagebrush (Artemisa californica) (Fig. 46). This sweet-smelling plant has medicinal properties and can be used as a topical pain salve.
Poza Canyon. As you walk through Poza Canyon, look for the lemonade berry shrub (Rhus integrifolia). This plant (Fig. 47) produces a sticky berry that can be placed in water to create a lemonade-flavored drink or sucked on to quench thirst.
Invasive plants. Santa Cruz Island was ranched from around 1850 to the 1980s. During this time, native vegetation was heavily impacted by grazing from sheep and rooting from domestic and feral pigs. Also, during this time, non-native species, such as European grasses and fennel were introduced to support the ranching and human life on the island. Other weeds were also accidentally or purposely introduced. Most of these introduced plants did not have natural predators, and some did well growing in the disturbed soil due to the livestock activity on the island. This created an ideal habitat for these invasive plant species to take over and crowd out native vegetation. Many efforts have been made and continue to be made to remove these invasive species and restore native habitat, although the impact of this period is still very noticeable today.
Around the Field Station. The large eucalyptus stands (Fig. 48) that surround the UC Field Station are an example of introduced species to Santa Cruz Island. These trees consume large amounts of water and can significantly alter valuable water flow on the island. Efforts are being made to remove some of these trees in order to restore water flows and former wetland habitat (such as the wetland by Prisoners Harbor).
Fauna of Santa Cruz Island
Island fox. Island foxes (Urocyon littoralis) (Fig. 49) can also be spotted around the UC Field Station. These little animals are the largest carnivore on SCI and are considered a keystone species. The island foxes almost faced extinction during the ranching era on the Channel Islands. During the ranching era, domestic pigs escaped and became feral throughout the island. Around the same time, bald eagles, which nested on the island, were being impacted by the chemical DDT (dichlorodiphenyltrichloroethane). The DDT, which was being disposed of in the ocean by industrial companies, was traveling up the food chain and into the bald eagle’s primarily marine diet. The chemical caused the bald eagle’s eggs to become too fragile and not viable. The bald eagles eventually left the Islands, and golden eagles moved in. The golden eagles were initially attracted to the feral piglets as a food source, but soon discovered the easy prey of the island foxes (which had previously had no natural predators). The golden eagles almost eliminated the fox population on the majority of the Channel Islands. Through human intervention, the golden eagles were removed from SCI and relocated. Feral pigs and other domestic ranch animals were also removed from the island. A captive breeding program restored the island fox population and bald eagles were once again reintroduced to the Channel Islands. The island fox population was able to rebound (Island Fox, n.d.).
Island scrub-jay. The island scrub-jay (Aphelocoma insularis) (Fig. 50) is endemic to Santa Cruz Island and found no place else in the world. Avid birders come from all over the world to see this bird. The island scrub-jays are larger and bluer than their mainland relatives. These personable birds are important seed dispersers for the oak trees on the Channel Islands. They are commonly seen at Prisoners Harbor, which is our Stop 1 (Island Scrub-Jay, n.d.).
PREHISTORIC CHERT USE ON EAST SANTA CRUZ ISLAND
During the late Holocene, Santa Cruz Island, the largest and most ecologically diverse of the eight Channel Islands, was home to substantial populations of Island Chumash people (Munns and Arnold, 2002). These Chumash communities, including those on Santa Cruz Island, were primarily coastal, sedentary, averaged 50–250 persons per village, and relied on a wide variety of nearshore and offshore marine resources and local and imported plant resources (Arnold, 1992, p. 66). The island’s prehistoric chronology is shown in Table 1.
|Period||Years before Present|
|Period||Years before Present|
East Santa Cruz Island Chert Sources
Near the El Montañon highlands—aka Montañon Ridge—a high volcanic range that divides the eastern tip of Santa Cruz Island from the rest of the landmass (Fig. 51), are more than a dozen large outcroppings of chert that were quarried on a small scale for thousands of years to make various implements (Arnold and Walsh, 2010, p. 117). Large outcroppings of chert-bearing strata of the Monterey Formation are exposed at the surface immediately west of this range, and smaller outcrops are found to the east (Munns and Arnold, 2002, p. 128). The most abundant and highest-quality chert sources are situated in and around El Montañon, the prominent ridgeline that separates the east end of Santa Cruz Island from the rest of the island (Perry and Jazwa, 2010, p. 180).
The Miocene-aged Monterey Formation, identified as marine biogenic shale, was mapped by Weaver and Nolf (1969) and Dibblee (2001b) and described by Weaver and Meyer (1969) and Dibblee (2001b). The Monterey Formation (Tm) shale member is buff to white, chalky and siliceous with thin-bedded chert, bentonite; contains foraminifers, moderately lithified; and as much as 1300 m in exposed thickness (Weaver and Meyer, 1969; Dibblee, 2001b). More detail is provided on the Monterey Formation here: the road log descriptions at Stops 9, 9B, 9C, 13, 13B; and see topical note section on “The Monterey Formation.”
Major chert source localities are restricted to the eastern end of the island, primarily to an area known as the “Contact Zone,” where blonde-to-brown cherts and shales of the Monterey Formation meet the massive Santa Cruz Island Volcanics (Fig. 51) (Arnold et al., 2001, p. 115). The cherts from these quarries are a visually distinctive variant of Monterey Formation chert (Arnold et al., 2001, p. 113). Commonly called “Santa Cruz Island Blonde chert” due to its light-brown color, this material is well-suited for flintknapping due to its high-silica content and blocky, rather than laminar structure (Arnold et al., 2001, p. 115; Perry and Jazwa, 2010, p. 180). Flintknapping is defined as forming stone implements by controlling the fracture of the objective piece (Andrefsky, 2005, p. 255).
In all, some 26 chert quarries have been documented on eastern Santa Cruz Island (Perry and Jazwa, 2010). This cluster of chert quarries was the only major source of medium- to high-grade chert near the coastline between Point Conception and San Diego, making it a valuable and controllable resource (Arnold and Walsh, 2010, p. 117). As well, the chert outcrops on eastern Santa Cruz Island were an important resource to the Chumash since chert resources are more limited on the other Channel Islands (Perry and Jazwa, 2010).
Exploitation and Utilization of Santa Cruz Island Chert Resources
While the Santa Cruz Island chert outcrops were used throughout the last ~10,000 years, there was significant variability in chert exploitation during this time span (Perry and Jazwa 2010). Scattered biface preforms and debitage at local sites show that early exploitation of the quarries focused on making large cores and bifacial tools (Arnold and Walsh, 2010, p. 117). A core is the nucleus or mass of rock that functions primarily as a source for detached pieces (Andrefsky, 2005, p. 254). A biface is a tool that has two surfaces (faces) that meet to form a single edge that circumscribes the tool (Andrefsky, 2005, p. 253). Preforms represent the stage of production of a lithic tool just prior to reaching a finished form (Andrefsky, 2005, p. 260). Debitage consists of detached pieces that are discarded during the reduction process (Andrefsky, 2005, p. 254).
During the middle Holocene, chert quarrying appears to have been dispersed and opportunistic, and was focused on the production of bifaces, drills, and other chert tools. Drills are tools that are used in a rotary motion to perforate materials. Prior to the advent of the microlithic industry, and extending into more recent times as a secondary use, the Santa Cruz Island chert also served as raw material for making small numbers of drills, which were used for drilling wood, larger shell artifacts, and other materials, as well as for bifacial tools (Arnold et al., 2001, p. 113).
In the late Middle period (see Table 1), flintknappers began to make modest numbers of small chert drills for perforating shell disk beads, and these have been found at several dozen coastal Chumash sites on both sides of the channel (Arnold and Walsh, 2010, p. 117; Munns and Arnold, 2002, p. 143).
By about A.D. 1150–1200, the Island Chumash appear to have secured access to the large chert quarries along the El Montañon Ridge and at nearby outcrops on the eastern portion of Santa Cruz Island (Fig. 51) (Munns and Arnold, 2002, p. 143). The blocky chert from these quarries is well suited for making microblades, and had been used for the preceding 200–300 years by both mainland and island residents for the low-intensity manufacture of microblades (Munns and Arnold, 2002, p. 143). Microblades are whole or fragmentary microliths (narrow and <50 mm long), produced from microblade cores, with no bits (Arnold et al., 2001, p. 114). Microliths are very small blades usually in geometric form that are used in composite tools.
Two types of microblades have been identified: (1) trapezoidal, and (2) triangular with dorsal retouch (TDR). The trapezoidal form dominated the late Middle period, whereas the TDR microblades were produced in significantly higher quantities during the Late period (Perry, 2004).
As the Transitional period progressed, the communities participating in microblade and shell bead manufacture became confined almost exclusively to the Northern Channel Islands. Specialists at sites within 8 km of the chert-bearing zone, including Chinese Harbor, Prisoners Harbor, a few sites on the eastern tip of Santa Cruz Island, and several quarries, manufactured millions of microblades over the next 500–600 years (Munns and Arnold, 2002, p. 143).
Shell Bead Manufacture
During the Late Holocene, chert quarrying was focused on microlith production associated with the manufacture of shell beads (Perry and Jazwa, 2010). Most of the microblades were transported to a number of specialized bead-making communities elsewhere on the island, as well as to Santa Rosa and San Miguel Islands (Munns and Arnold, 2002, p. 143).
From the 1100s to the early 1800s, tremendous quantities of chert microdrills and Olivella beads were manufactured on the islands on a sustained basis (Arnold and Walsh, 2010, p. 117). Microdrills are modified microblades and include some or all of a retouched bit (Arnold et al., 2001, p. 114).
Chert microdrills were used to drill callus cup beads from the purple dwarf olive (Olivella biplicata, a small sea snail) shell. Olivella callus beads were used historically as a standard of value and reportedly functioned as currency during the Late period (Arnold, 1992, p. 73).
Many archaeologists have described the islands as the exclusive “mint” for the shell bead money, which was used throughout southern California (Arnold and Walsh, 2010, p. 119). The Santa Cruz Island quarry zone and nearby villages represent the most intensive, large-scale microblade industry known in the Americas. The roughly 20 specialized bead-making villages on Santa Cruz, Santa Rosa, and San Miguel Islands constituted perhaps the most extensive pre-Columbian shell-working industry north of Mexico, even surpassing the scale of shell crafts associated with the great Mississippian cultures (Arnold and Walsh, 2010, p. 119).
Large-scale microlith and bead-manufacturing specialization persisted throughout the Late period and well into the Historic era. Islanders were the primary suppliers of beads for the entire Chumash sphere, and a number of their mainland neighbors such as the Gabrielino/Tongva. Even after glass beads and iron needles introduced by the Spanish began to circulate in notable numbers (~A.D. 1782) in both mainland and island contexts, islanders continued to make very large quantities of shell beads, some with chert microdrills and some with needles (Munns and Arnold, 2002, p. 143).
The field-trip leaders would like to thank Marc Kamerling, Scott Minor, Dan Muhs, Richard Heermance, Geoff Gallant, and Mathew Davis for their thoughtful reviews and edits of this manuscript. We also want to recognize Lindsey Hronek’s careful and colorful drafting of most of the figures. We also wish to thank several individuals and organizations that made this trip possible: Drs. Lyndal Laughrin and Jay S. Reti, plus Brian Guerrero of the Santa Cruz Island Reserve Field Station (UC Santa Barbara Natural Reserve System) for their very helpful advice and assistance in organizing the field-trip logistics. We also thank The Nature Conservancy and Channel Islands National Park for allowing access to the island, and the Santa Cruz Island Reserve Field Station, which provided accommodations and vehicles for our visit.