Refined assessment of the paleoceanographic and tectonic influences on the deposition of the Monterey Formation in California

Famous for its petroleum potential and extensive deposits of fined-grained biosiliceous rocks, the Monterey Formation has been extensively mapped and studied throughout western California (Fig. 1). The Monterey Formation and similar biosiliceous rocks span the interval from the late early Miocene through the late Miocene, documenting both the climatic and tectonic history of coastal California.

Abundant marine diatoms characterize the original component of most of the Monterey Formation, making diatom biostratigraphy a very useful tool for correlating and dating stratigraphic sections in California and in onshore and offshore areas around the North Pacific Ocean to Japan (Ingle, 1981; Koizumi, 1973). However, diagenesis in siliceous mudstone caused by burial and exposure to temperatures in excess of 50 °C causes destruction of most diatom frustules during the conversion of opal A to opal CT. Similarly, opal CT is converted to quartz (chert) at temperatures between 65 °C and 80 °C (Keller and Isaacs, 1985). In such cases, diatoms may be preserved if they have been encapsulated in calcium carbonate concretions or dolomite that formed early in the diagenetic process (White et al., 1992).

Pioneering papers by Koizumi (1973) and Schrader (1973) on the Miocene to recent diatom biostratigraphy of Deep Sea Drilling Project cores from the North Pacific Ocean provided the framework for the development of Barron’s (1976b) diatom biostratigraphy for the Monterey Formation. Subsequent refinement of diatom biostratigraphy followed during the next 20 yr, including studies by Barron (1980, 1981), Akiba (1986), Barron (1986a), and Barron and Gladenkov (1995), with a synthesis published by Yanagisawa and Akiba (1998). Unlike calcareous nannofossils and planktonic foraminifers, diverse diatom assemblages are found in colder waters, where they often are the main tools for establishing biochronology. In particular, planktonic diatoms have proven to be most useful for biostratigraphy, especially species of the cold-water genera Denticulopsis and Neodenticula, along with diverse forms of the genus Thalassiosira.

Figure 2 compares the California diatom zonation of Barron (1986a, this paper) with the North Pacific diatom zonation of Yanagisawa and Akiba (1998) and the calcareous nannofossil zonation of Okada and Bukry (1980) with the paleomagnetic time scale of Gradstein et al. (2012). Correlation of the calcareous nannofossil zones is after Raffi et al. (2006), whereas correlation of the diatom zones is after Yanagisawa and Akiba (1998) and Barron and Isaacs (2001).

Using this refined biostratigraphy and published studies on paleoceanography and tectonics over the past 33 yr, the purpose of this paper is to update the observations made by Barron (1986b) in order to provide an assessment of the imprint of global paleoclimatic and paleoceanographic events and regional tectonic events on deposition of the Monterey Formation and post-Monterey biosiliceous rocks in California.

Figure 3 summarizes the correlation of key stratigraphic sections in California (Fig. 1) to diatom biostratigraphy and to the geologic time scale of Gradstein et al. (2012) (Fig. 2). Biostratigraphic studies used in compiling this figure include: Upper Newport Bay (Ingle, 1971; Warren, 1972; Barron, 1976b; Barron and Keller, 1983); Santa Monica Mountains (Blake, 1991; Barron, 1976a); Naples coastal bluffs (Barron, 1986a; Arends and Blake, 1986; Hornafius, 1994; Föllmi et al., 2005); San Joaquin Valley (Graham and Williams, 1985; Mosher et al., 2013; Hosford Scheirer and Magoon, 2007); Salinas Valley (Durham, 1974; Poore et al., 1981; Baldauf and Barron, 1982; Stanley et al., 2017; Barron, 2021, personal observ.); and the Monterey–Santa Cruz area (Barron, 1976b; Clark, 1981; Dumont et al., 1986; Clark et al., 1997; Powell et al., 2007). The stratigraphic columns for the Salinas Valley and San Joaquin Valley are generalized. Correlation of California benthic foraminiferal stages is based on their correlation with calcareous nannofossil and diatom zones in the Upper Newport Bay section with reference to Crouch and Bukry (1979).

The base of Monterey Formation is defined by reduced terrigenous input (Isaacs, 1985, 2001). Typically, the base is ca. 16 Ma or near the Relizian-Luisian benthic foraminiferal stage boundary, but it can be older near the center of basins or younger and unconformable with underlying rocks on basin margins. For example, the base of the Monterey Formation ranges from ca. 18 Ma in the Naples coastal bluffs section west of Santa Barbara (Figs. 1 and 3) to ca. 16 Ma in the Los Angeles Basin, the Santa Maria Basin, the San Joaquin Valley, the Salinas Valley, and the Santa Cruz area (Figs. 1 and 3). However, in the Santa Monica Mountains and Puente Hills of southern California, the base of Monterey-correlative rocks (the Modelo and Puente Formations) is ca. 13–12 Ma, unconformably overlying clastic rocks of the Topanga Formation (Figs. 1 and 3; Blake, 1991).

The base of the Monterey Formation has also been tied to the onset of diatom deposition in California and elsewhere in the North Pacific Ocean (Bramlette, 1946; Ingle, 1981; Keller and Barron, 1983). However, Isaacs (1985) argued that study of underlying clay-rich strata in the Rincon Shale at Naples showed that biosilica influx (when corrected for the effects of detrital dilution) in that unit equaled average rates in the Monterey Formation (2 mg/cm²/yr) and considerably exceeded middle Miocene rates. Isaacs (1985) suggested that diatom productivity patterns previously thought to be exclusively characteristic of the Monterey Formation began at least by the beginning of the Miocene (ca. 23 Ma). In fact, studies by the author of 190 samples from southern California Borderland seafloor outcrops taken by the U.S. Geological Survey (USGS) during the late 1970s revealed five diatom-rich early Miocene samples that were older than the Crucidenticula sawamurae zone (18.2 Ma; Fig. 2). Again, the distinctive feature about the Monterey Formation is an unusually slow influx of terrigenous debris that would otherwise have diluted the biogenous components. In other words, according to Isaacs (1985, 2001), the principal cause of the lower formation boundary was a sharp reduction in the rate of influx of terrigenous debris.

The fine-grained nature of the Monterey Formation has been attributed to both regional tectonics and sea-level fall (Ingle, 1981; Barron, 1986b; Behl, 1999). The development of a transform margin ca. 18 Ma (Atwater, 1998) created a series of parallel offshore basins, increasingly separated from onshore clastic deposition by structural highs (Ingle, 1981). Rising sea levels during the middle Miocene (Barron, 1986a; Behl, 1999) caused further restriction of clastic deposition to nearshore basins, paving the way for the deposition of fine-grained sediments in more offshore basins.

Figure 4 reveals that a major rise in sea level ca. 18 Ma coincides with the base of the Monterey Formation in the Naples section. The 16 Ma base of the fine-grained rocks of the Monterey Formation in many other sections in California (Fig. 4), however, does not correlate with a major rise in sea level according to the sea-level curve of Kominz et al. (2008), which was developed on the New Jersey margin.

Although Kominz et al. (2008) did not recognize a fall in sea level at the ca. 13 Ma time of the onset of deposition of terrigenous-rich sediments of the Santa Margarita Sandstone in the Salinas Valley (Baldauf and Barron, 1982), the more recent reconstruction of Miller et al. (2020) did recognize this event.

Compressed intervals are recorded in the Newport (ca. 12.3–10.0 Ma) and Naples (ca. 13.0–10.6 Ma; Föllmi et al., 2005; Blake, this volume) sections. John et al. (2002) documented a compressed interval between ca. 12.7 and 10.8 Ma in the section at El Capitan State Beach, immediately west of the Naples section, but compressed intervals are also observed in sections in the Santa Maria Basin (White et al., 1992), where they are associated with phosphatic sediments, as they are at Naples and El Capitan State Beach.

In the El Capitan section west of Naples, John et al. (2002) noted that the phosphate beds show different indications of sediment reworking, including fracturing and reworking of phosphate laminae, heterogeneous mixtures of phosphate particles, the presence of different phosphate generations around nodules, and the presence of erosional structures, an observation supported in the Naples section by Föllmi et al. (2005). This later paper argued that seismic activity was instrumental in shaping the sediments of the Monterey Formation. It proposed that in situ deformation of sediments facilitated the erosion and reworking of preexisting sediments and episodically triggered slides and gravity flows. Föllmi et al. (2005) suggested that the Naples section was deposited on a low-gradient slope, perhaps in a base of slope or a basin apron position, probably downslope from an offshore submarine ridge. They argued that the condensed interval between 13 and 10 Ma was associated with low mass accumulation and silica dissolution, possibly caused by eustatically controlled sediment starvation. This hypothesis is supported by the relatively high sea level reported by Kominz et al. (2008) during this interval.

Deposition of diatom-rich sediments during the 11.5–10.0 Ma interval is documented in the Palos Colorados Mine (also known as the Grefco Mine), southeast of Lompoc (Knott et al., this volume; Barron, 1986a), the Toro Road section east of Monterey (Barron, 1976b, 1986a), and in the Chico Martinez Creek section of the southern San Joaquin Valley (Mosher et al., 2013). Barron et al. (2002) revealed that a dramatic threefold decline in biogenic opal abundance in offshore Ocean Drilling Program (ODP) sections began at 11.5 Ma, arguing for a steepening of the offshore-onshore productivity gradient. Interestingly, this compressed interval coincides with a period of higher sea level according to Kominz et al. (2008), possibly suggesting high sediment accumulation rates in inboard sections closer to the strand line and sediment starvation in offshore areas.

A major fall in sea level at ca. 10 Ma is associated with the base of the Santa Margarita Sandstone in the Toro Road section east of Monterey (Barron, 1976b), and it may actually date the base of the Santa Margarita Sandstone in the Santa Cruz area, where an unconformity separates the sandstone for the underlying Monterey Formation (R. Stanley, 2019, personal commun.). Subzone d of the Denticulopsis lauta–Denticulopsis hustedtii zone (also referred to as the Denticulopsis dimorpha zone) spans the interval between 10.0 and 9.3 Ma (Fig. 2) and commonly marks the onset of deposition of widespread diatomaceous rocks in the Los Angeles and Ventura Basins, attesting to rising sea level after 10 Ma (Barron, 2021, personal observ.).

A new study by Miller et al. (2020) proposed that eustatic sea level rose ~20 m between ca. 8.0 and 7.5 Ma, before remaining near modern levels for the rest of the late Miocene (Fig. 4). This conflicts with the sea-level curve of Haq et al. (1987), which indicates a 30 m drop in sea level centered on 6.5 Ma, a correlation made by Barron (1986b) with the top of the Monterey Formation.

Graham and Williams (1985) pointed out that the change from a convergent to a transform Pacific–North American plate margin imposed a new structural style on coastal California. Transform faults parallel to the coastline created series of onshore-offshore basins wherein terrigenous debris was trapped in the inner basins (Ingle, 1981), allowing deposition of fine-grained, organic-rich sediments in outer basins. Accelerated subsidence between 16 and 14 Ma occurred in the western San Joaquin Basin according to Graham and Williams (1985).

Beginning at ca. 18 Ma, the California margin between the southern California Continental Borderland and the Mendocino triple junction corresponded to the oblique transtensional portion of the Pacific–North American plate boundary (Atwater, 1998). The northern terminus of this boundary migrated northwestward up the coast, attaching to the Pacific plate at the Mendocino triple junction between 19 and 12 Ma.

Rotation of the Western Transverse Ranges dominated the Miocene tectonic history of southern California. Paleomagnetic data from the Ventura region and from both sides of the Santa Barbara Channel reveal that rocks older than ca. 16 Ma have been rotated 80°–110° clockwise, while younger rocks show progressively lesser amounts of rotation (Luyendyk et al., 1980; Hornafius, 1985; Luyendyk, 1991), whereas most nearby rocks outside the Western Transverse Ranges show much less overall rotation. The widely accepted explanation of these results is that the Transverse Ranges block rotated 90°–110° as a more or less coherent piece since the early Miocene.

Ingersoll and Rumelhart (1999) proposed a three-stage model for evolution of the Los Angeles Basin and vicinity within the evolving transform-fault system: (1) transrotation (18–12 Ma), (2) transtension (12–6 Ma), and (3) transpression (6–0 Ma). Ingersoll and Rumelhart (1999) argued that the timing of these stages correlated with microplate-capture events, which occurred during conversion from a convergent margin to a transform margin.

Rumelhart and Ingersoll (1997) employed subsidence and provenance analyses in proposing a paleogeographic and paleotectonic reconstruction of southern California. Transrotation of the Western Transverse Ranges, which began at 16 Ma, led to extension and thermal subsidence of the Los Angeles Basin area (Fig. 5). A second pulse of extension and thermal subsidence began at 12 Ma, when motion began along the San Gabriel fault (Fig. 1), Transtension in the Los Angeles Basin area and deposition of the Puente, Tarzana, Simi, and Piru fan systems coincided with right slip of 60–70 km along the San Gabriel fault. Transform motion was transferred to the San Andreas fault at 6 Ma, causing transpression to dominate the Los Angeles Basin since 6 Ma and inducing “rapid uplift, flexural subsidence due to tectonic loading, and rapid sedimentary filling” (Rumelhart and Ingersoll, 1997, p. 885) (Fig. 5).

In southern California, the onset of deposition of the Monterey-like Modelo and Puente Formations in the Santa Monica Mountains and Puente Hills occurred at 13–12 Ma above unconformities with the underlying clastic-rich Topanga Formation (Blake, 1991). Rumelhart and Ingersoll’s (1997) 12 Ma onset of extension and thermal subsidence in the Los Angeles Basin may have led to onset of deposition of the Modelo and Puente Formations.

Figure 6 compares the benthic foraminiferal oxygen isotope curve of Zachos et al. (2001, 2008) with the representative sections of the Monterey Formation in California. Significant climatic and paleoceanographic events that appear to have affected deposition of the Monterey Formation and overlying biosiliceous rocks are included on the figure.

The Miocene climatic optimum refers to the interval of time between 16.9 and 14.7 Ma when high global mean annual temperature was accompanied by a relatively low global CO₂ concentration (Holbourn et al., 2007, this volume; Sosdian et al., 2020). The coincident “Monterey excursion” of Vincent and Berger (1985) was a long-lasting positive carbon isotope excursion that ended ca. 13.5 Ma, ~400 k.y. after major expansion of the Antarctic ice sheet (Holbourn et al., 2007, this volume). The apparent covariance between δ¹³C and δ¹⁸O gave support to the hypothesis that periodic increased burial of organic carbon drove atmospheric CO₂ drawdown, spurring Miocene global cooling (Vincent and Berger, 1985; Woodruff and Savin, 1991; Flower and Kennett, 1993, 1994; Sosdian et al., 2020).

The Relizian and Luisian benthic foraminiferal stages represent the Miocene climatic optimum in California, ranging from the upper Crucidenticula sawamurae through Denticulopsis hyalina diatom zones (Fig. 2). This interval in the Monterey Formation typically contains tropical planktonic foraminifers and calcareous nannofossils (Poore et al., 1981; Barron and Keller, 1983; Flower and Kennett, 1993, 1994; Föllmi et al., 2005), but diatom assemblages are dominated by North Pacific taxa with relatively few tropical species (Barron and Keller, 1983; Barron, 2003).

According to the benthic δ¹⁸O data of Zachos et al. (2001, 2008), the Miocene climatic optimum was followed by the middle Miocene climate transition (ca. 14.7–13.8 Ma; Holbourn et al., this volume) and then by progressive high-latitude cooling between ca. 13.8 and 8.0 Ma (Fig. 6). Increasing diatom deposition and decreasing calcium carbonate deposition in the Monterey Formation and similar sediments in the North Pacific Ocean (Barron and Keller, 1983; Barron, 1986b) accompanied this cooling.

Past literature (Bramlette, 1946; Ingle, 1981; Keller and Barron, 1983) has equated the base of the Monterey Formation with the onset of widespread diatom blooms along the coast of California. Following arguments made by Berger (1970), who stated that basin-basin fractionation of nutrients explains why sediments of the North Pacific Ocean are diatom rich, while those of the North Atlantic Ocean are diatom poor, Keller and Barron (1983) proposed a silica shift from the North Atlantic Ocean to the North Pacific Ocean at ca. 16–15 Ma, coincident with the base of the Monterey Formation. Diatom-rich deposits, however, were still common along the New Jersey margin as late as 13 Ma (Barron et al., 2013).

Isaacs (2001) pointed out that the Monterey Formation is often viewed as a straightforward paleoceanographic record deposited on the floor of steep-sided restricted basin like the modern southern California Continental Borderland basins. She stated that the late early and middle Miocene (18–11 Ma) midbathyal slope on the ocean margin was characterized by moderate biosilica accumulation. Isaacs (2001) suggested that the 11–5 Ma interval was marked by regionally high biosilica accumulation in coastal Californian basins, in comparison with the San Joaquin Basin, where biosilica accumulation was high between 17 and 5 Ma. According to Hosford Scheirer and Magoon (2007), the onset of the biosiliceous facies of the Monterey Formation in the San Joaquin Valley occurred between 13.5 Ma (the McDonald and Fruitvale Shale) and 12 Ma (the McLure Shale). Rick Behl (2019, personal commun.), however, argues that the base of the biosiliceous Antelope Shale is closer to 11.5 Ma in the Belridge Field of the southern San Joaquin Valley (Fig. 3).

So, does this apparent 13.5–11.5 Ma increase in diatom deposition in the Monterey Formation correlate with increased diatom productivity elsewhere in the Pacific Ocean and a shift in silica from the North Atlantic Ocean? It is useful to review the detailed records of Miocene biogenic opal productivity compiled in the equatorial Pacific Ocean by Lyle and Baldauf (2015). Lyle and Baldauf (2015) documented a middle Miocene regime older than 13.2 Ma marked by low diatom numbers and a few short-lived productivity intervals, followed by a carbonate crash regime. The carbonate crash regime began with an older substage (13.2–10.2 Ma) characterized by relatively high abundances of productivity-related diatoms and common productivity-related depositional intervals interspersed with CaCO₃ dissolution intervals. A younger carbonate crash substage (10.2–8.0 Ma) followed, marked by weaker brief productivity intervals, high CaCO₃ dissolution, and moderately high numbers of upwelling diatom species (Fig. 6). After this interval, a biogenic bloom regime (8.0–4.5 Ma) occurred, with extended periods of high opal and CaCO₃ deposition and a maximum between 7.0 and 6.4 Ma. Finally, a Pliocene–modern regime (4.5–0 Ma) arose with lower productivity and high cyclic CaCO₃ dissolution (Fig. 6).

The carbonate crash regime included an older (13.2–10.2 Ma) low-CaCO₃ interval found on both sides of the Isthmus of Panama (Roth et al., 2000) and a later (10.2–8.0 Ma) carbonate crash CaCO₃ minimum found only on the Pacific side (Lyle et al., 1995; Nathan and Leckie, 2009). The onset of the carbonate crash at 13.2 Ma (Lyle and Baldauf, 2015) coincided with decline of diatoms in the New Jersey offshore record (Barron et al., 2013) and an increase of diatom deposition in onshore sections of northeast Japan (Koizumi et al., 2009).

The modern Pacific Ocean is enriched in nutrients compared to the North Atlantic Ocean primarily due to the production and flow path of North Atlantic Deep Water (NADW).

Cd/Ca records from Delaney and Boyle (1987) showed that the distribution of nutrients within the Atlantic and Pacific Oceans started to diverge significantly at 12.5 Ma. The emergence of a volcanic arc south of Central America and collision with northern Colombia (Coates et al., 2003) between 12.8 and 7.1 Ma were likely important events in isolating equatorial Pacific Ocean waters from those of the North Atlantic, and therefore leading to increasing nutrients in the Pacific Ocean, which were favorable for diatom production. Model simulations by Butzin et al. (2011) of the early and middle Miocene support the conclusions by Woodruff and Savin (1989) and Wright et al. (1992) that deep-water formation in the North Atlantic Ocean was absent or weak prior to 11 Ma. It would seem that the late middle Miocene (ca. 13–11 Ma) increase of diatom deposition in the North Pacific Ocean coincided with increased production of NADW in the North Atlantic Ocean, which enhanced North Atlantic–North Pacific basin-basin fractionation.

As pointed out by Barron (1998), there is evidence that opal accumulation rates actually increased in onshore sections of southern California beginning at ca. 8.0 Ma (age updated), coincident with the lower part of the biogenic bloom regime of Lyle and Baldauf (2015) in the equatorial Pacific Ocean (Fig. 6). In the Naples section, this interval coincides with the CaCO₃-poor, informally named clayey-siliceous member of the Monterey Formation of Isaacs (1981, 1984) (Fig. 3) and marks an up-section doubling of silica accumulation rates (Isaacs, 1984, 1985). Along the Santa Cruz coast, this interval is marked by the thick sequence of the biosiliceous Santa Cruz Mudstone (Fig. 3). At Lompoc, the most pure (mineable) deposits of the Manville Quarry occur between ca. 8.4 and 7.4 Ma, implying enhanced diatom production. Barron et al. (2002) documented a decline in opal abundance in ODP sites from offshore California, which appears to have been related to an increasing onshore-offshore productivity gradient.

In a synthesis of alkenone sea-surface temperature studies, Herbert et al. (2016) documented a major cooling in both hemispheres occurring between ca. 7.5 and 6.5 Ma. This event coincided with the first significant ice in Greenland (Larsen et al., 1994) and South America (Mercer and Sutter, 1982), the base of the glacial Yakataga Formation in coastal southern Alaska (Eyles et al., 1991), a stepped decrease in carbon isotope records, suggesting a drawdown of atmospheric CO₂ (Bradshaw et al., 2012), and the onset of increased seasonality and aridity in the Northern Hemisphere (Diester-Haass et al., 2006). Sea-surface temperatures fell to near-modern values during this 7–5.4 Ma interval, coincident with near-modern sea levels (Fig. 4). Presumably, a steepening of latitudinal thermal gradients during this latest Miocene time would have led to an increased wind-derived coastal upwelling along the California coast, resulting in enhanced accumulation rates of biogenic opal (predominantly diatoms).

As noted by Isaacs (1981, 1985) and Barron (1986a, 1986b), the top of the Monterey Formation, which is marked by the contact with overlying, more terrigenous-rich sediments (the Capistrano, Sisquoc, Poncho Rico, and Purisima Formations), is typically abrupt in coastal California and often is marked with an unconformity (Fig. 7). Detailed study of the Monterey-Sisquoc boundary in the Santa Maria Basin by Barron and Ramirez (1992) revealed considerable variability in nature of the contact. In some places, this boundary coincides with the opal CT/opal A boundary. Isaacs (1985) pointed out that diagenesis is hindered by greater terrigenous-rich components, implying that the less clastic-rich Monterey Formation would be more likely to be converted to opal CT. Thus, the opal A/opal CT boundary commonly coincides with the top of the Monterey Formation.

Graham and Williams (1985), however, pointed out that the upper transition to clastic-rich sediments is time-transgressive in the San Joaquin Valley, with the base of the clastic-rich Etchegoin Formation ranging from near the Miocene-Pliocene boundary in the basin center to near the base of the Thalassiosira antiqua zone (8.6 Ma) elsewhere. Similarly, clastic deposition began in the Modelo Formation of the Santa Monica Mountains (Figs. 1 and 3) at ca. 9 Ma (Blake, 1991; Rumelhart and Ingersoll, 1997), well below the top of the Monterey Formation in coastal California.

In contrast to the argument of Barron (1986b), which stated that this 7 Ma contact coincided with a 30 m drop in sea level centered on 6.5 Ma, as shown by Haq et al. (1987), a new late Miocene sea-level synthesis of Miller et al. (2020) did not suggest any appreciable fall in sea level during this interval (Fig. 7). Nevertheless, major global cooling and the initiation of glaciation in both the Northern and Southern Hemispheres (Herbert et al., 2016) would seem to call for a fall in global sea level near the end (ca. 7 Ma) of Monterey Formation deposition.

Evidence is unclear, however, whether this 7 Ma event coincided with a major tectonic event in coastal California. In southern California, Rumelhart and Ingersoll (1997) argued that transform motion was transferred from the San Gabriel fault to the San Andreas fault at ca. 6 Ma, noting that transpression has dominated the Los Angeles Basin since 6 Ma, including rapid uplift, flexural subsidence due to tectonic loading, and rapid sedimentary filling.

Could this tectonic event actually have occurred at ca. 7 Ma rather than at 6 Ma? McCrory et al.’s (1995) review of Neogene deposition and structure in the Santa Maria Basin pointed out that all sections with detailed geohistories record a complicated latest Miocene episode at the Monterey-Sisquoc boundary (ca. 7 Ma, age updated) characterized by either deepening or shoaling. The time-transgressive nature of the opal A/opal CT boundary in the Santa Maria Basin (Barron and Ramirez, 1992) would appear to reflect basin deepening, leading to higher burial temperatures and conversion to opal CT, versus basin shoaling, leading to preservation of opal A.

Latest Miocene (younger than 6.6 Ma) marine diatoms occur in the basal St. George Formation at Crescent City (Fig. 1) overlying a paleosol that developed on the Mesozoic Franciscan Complex (Robinson et al., 2001). Aalto (2006) attributed this horizon to rapid submergence of northwestern California to bathyal depths. Similarly, an unconformity spanning the interval from ca. 7.8 to 5.9 Ma is present in the Pullen Formation at Centerville Beach in the Eel River Basin (McCrory, 1990).

In summary, there is evidence in both the Santa Maria Basin and in northernmost California at Crescent City that a tectonic event occurred at ca. 7 Ma coincident with the top of the Monterey Formation. Whether this tectonic event occurred throughout California and possibly was associated with the San Andreas fault is currently unclear.

The early Pliocene interval from 4.6 to ca. 3.5 Ma is notable by greatly reduced opal sedimentation offshore California (Barron, 1981, 1998; Barron et al., 2002; Fig. 2) and is correlative with the onset of the Pliocene–modern regime (4.5–0 Ma) of Lyle and Baldauf (2015) in the equatorial Pacific Ocean, a period of lower opal productivity and high cyclic CaCO₃ dissolution (Fig. 6). Kwiek and Ravelo (1999) argued that North Pacific Intermediate Water (NPIW) ventilation from the North Pacific Ocean was enhanced during this time, resulting in a deeper nutrient maximum in intermediate waters of California margin ODP Sites 1014 and 1018. Onshore, this interval also coincided with terrigenous-rich sedimentation that completely masks any diatoms (Barron, 1992). DeMets and Merkouriev (2016) concluded that the Pliocene orogeny in western California began after 5.2 Ma, when the direction of Sierra Nevada/Great Valley–Pacific plate motion rotated clockwise of the local N41°W azimuth of the San Andreas fault.

In spite of its shortcomings in diagenetically altered rocks, diatom biostratigraphy offers a major tool for establishing the biochronology of the Monterey Formation and post-Monterey biosiliceous rocks in California. A robust correlation of both calcareous nannofossil and diatom biostratigraphy to the paleomagnetic time scale, which has been refined during the past 30 yr, provides the basis for assessing the imprint of global paleoclimatic and paleoceanographic events and regional tectonic events on the deposition of these Miocene to lowermost Pliocene rocks. Key stratigraphic sections from Upper Newport Bay, the type Modelo section of the Santa Monica Mountains, the Naples coastal bluffs section west of Santa Barbara, and the Monterey/Salinas area are correlated with generalized sections from the San Joaquin and Salinas Valley to provide a framework for assessing these paleoclimatic and paleoceanographic imprints on deposition. The abrupt decline in terrigenous deposition that marks the base of Monterey Formation generally occurred at 18–16 Ma, and it coincided with both a fall in global sea level (18 Ma) and tectonic deepening associated with the onset of transform faulting along the plate margin between the North American and Pacific plates. The Miocene climatic optimum (16.9–14.7 Ma) coincided with the presence of tropical calcareous microfossil assemblages in the lower Monterey Formation; however, diatom assemblages are largely provincial in character. Diatom deposition did not begin abruptly at the base of the Monterey Formation but existed earlier in offshore basins of southern California. A silica shift from the North Atlantic to the North Pacific Ocean did not occur at 16–15 Ma as argued by Keller and Barron (1983) but occurred later and more gradually between ca. 13 and 11 Ma, as the formation of North Atlantic Deep Water became established and deep-water flow across the Central American Seaway became restricted. Diatom deposition was enhanced in Monterey sections at ca. 11.5 Ma, but this interval is obscured in many coastal California sections (e.g., Newport and Naples) by the presence of a compressed interval, which seems to have been caused by sediment starvation in offshore basins during a period of rising sea level. Increasing latitudinal thermal gradients between ca. 13 and 8 Ma led to enhanced upwelling and deposition of diatoms along the California coast. Similar to the equatorial Pacific Ocean, coastal California experienced a biogenic bloom between ca. 8 and 5 Ma, as evidenced by increasing diatom accumulation rates. The expression of this event, however, is complicated by local tectonic events that enhanced terrigenous deposition, which can mask the diatom component. An unconformity or hiatus centered on 7 Ma is common in many coastal sections, where it typically separates the Monterey Formation from more terrigenous biosiliceous sediments of the Capistrano, Sisquoc, Poncho Rico, and Purisima Formations. A 7 Ma lithological break is not apparent in the Newport or Waddell Bluff (Santa Cruz coast) sections, suggesting that these sections were deposited in the deeper parts of basins. A new sea-level curve for the latest Miocene by Miller et al. (2020), however, does not indicate a fall in global sea level at 7 Ma, suggesting that something else was likely responsible for this 7 Ma break. Local and/or regional tectonics may have been responsible, based on evidence that both shoaling and deepening occurred coincident with the Monterey-Sisquoc boundary.

This paper benefited from discussions with Gregg Blake, Richard Behl, and Richard Stanley. Kris MacDougall provided a very helpful review for the U.S. Geological Survey (USGS), and the manuscript was improved through the reviews of two anonymous reviewers for the Geological Society of America. The USGS Land Change Science Program and Climate Research & Development Program and the Geology, Energy, Geophysics, and Mineralogy Science Center provided nonsalary support for this study.

In a historical context, I am extremely indebted to my Ph.D. dissertation advisors at the University of California Los Angeles, Helen Tappan and Alfred R. Loeblich Jr., who encouraged and supported my wish to study the diatoms of the Monterey Formation in California in 1971. I also acknowledge James C. Ingle Jr. and Robert E. Garrison for their and their students’ collaboration on Monterey studies over the past 40+ yr.