High-Resolution Stable Isotope Paleotopography of the John Day Region, Oregon, United States

Introduction

The John Day region lies in the High Desert of central Oregon with the Cascades Range to the west and the Blue Mountains province to the east (Figures 1A,B). The region, denoted here as the extent of the eastern and western facies (Figure 1B), is known for its near-continuous, fossil-rich terrestrial sedimentation, making it one of the planet’s most complete records of Cenozoic environments and climate. Over one hundred years of extensive study has built a foundation ripe for disentangling the co-evolution of local climate, landscapes, and life (Bestland and Retallack, 1994; Clark, 1989; Dillhoff et al., 2009; Merriam and Sinclair, 1906; Retallack et al., 2004; Robinson et al., 1984; Robson et al., 2019; Samuels et al., 2015; Sinclair, 1905; Swanson and Robinson, 1968). In addition to providing an ideal case study for paleo-biogeographical change (Kohn and Fremd, 2007; Retallack, 2004), John Day proxy data have been used to unravel the tectonic history of the Cascades, which serve as a barrier between Oregon’s High Desert and moisture-laden winter westerly winds originating in the Pacific (Bershaw et al., 2019; Kohn et al., 2002; Takeuchi et al., 2010). Due to the strong rainshadow cast by the Cascades today (Figure 1C), the range’s uplift history is considered an important driver of Cenozoic precipitation and vegetation change in the John Day region and across the northwestern United States (Retallack, 2004; Takeuchi and Larson, 2005; Kohn and Fremd, 2007; Takeuchi et al., 2010).

FIGURE 1. (A) Surface elevation (km) map of study region. Light blue represents ocean, diamond symbols are authigenic clay sample locations. Inset indicates the location of the study area within the United States. Elevation data collected with the elevatr package in R. (B) Approximate extent of the western (light blue polygon) and eastern (dark blue polygon) facies. Eastern facies follows the extent of the Blue Mountains region. Polygons based on Albright et al. (2008); Graham (2014); Robinson et al. (1984); Takeuchi (2007). Red box denotes the areal extent of maps in Figure 3. (C) Annual precipitation map (mm/yr). White area receives less than 400 mm of precipitation per year and characterizes the Cascades rainshadow. Data from the PRISM Climate Group (2012). (D) Aboveground biomass (kg/m²; live, dry). White area denotes no or negligible aboveground biomass. Data from Kellndorfer et al. (2012).

However, the Cascades Range is not the only topographic feature of consequence in the John Day region. The Blue Mountains province, which accreted in the Mesozoic (Dickinson and Thayer, 1978; Dickinson, 1979; Dickinson, 2004; Schwartz et al., 2010; LaMaskin et al., 2015), contains the eastern facies of deposition in the John Day region, separating it from the western facies in the adjacent high plains (Robinson et al., 1984) (Figure 1B). Today, the Blue Mountains create local gradients in precipitation and vegetation, with higher precipitation and a denser canopy in the mountain slopes (eastern facies) compared to the surrounding plains (western facies; Figures 1C,D). This significant local variability in hydroclimate and ecology may confound regional-scale interpretations of paleo archives if similar gradients existed in the past. For example, spatial variability can be misinterpreted as temporal variability if outcrop or sampling locations shift along these gradients with age. Similarly, changes in environmental parameters due to changes in local high elevation can be mis-attributed to the height of the Cascades if local topography is not accounted for.

There is some evidence that precipitation and vegetation gradients in the Blue Mountains have been present as far back as the Eocene. Near the western tip of the eastern facies Bestland and Retallack (1994) identified two different plant assemblages in the same stratigraphic horizon that point to distinct elevation regimes, perhaps related to Blue Mountain topography or local stratovolcanoes (White and Robinson, 1992). The floras are interpreted to reflect a wetter, more densely vegetated tropical lowland forest and an adjacent temperate, early successional highland forest (Bestland and Retallack, 1994; Bestland et al., 2002). In contrast, wetter conditions with denser vegetation appear in the highlands today while the lowlands are drier and sparsely vegetated (Figures 1C,D). Thus, while ecological gradients surrounding the Blue Mountains may be long-lived, their magnitude and even their direction may have changed with time.

In this study we present spatial and temporal oxygen isotope data from authigenic clay minerals, a proxy for past precipitation ${\delta }^{18}O$, to constrain the spatial expression of Blue Mountains topography and its effect on regional precipitation patterns. Isotopes of authigenic clays have been widely applied to studying tectonic and climatic change in Cenozoic western North America (Takeuchi and Larson, 2005; Mulch et al., 2006; Sjostrom et al., 2006; Mix et al., 2011, Mix et al., 2016, Mix et al., 2019; Chamberlain et al., 2012; Mix and Chamberlain, 2014) and they are particularly useful in places like John Day where soil carbonates—a more commonly used paleo-${\delta }^{18}O$ proxy—are not continuously present through the sedimentary record (Bestland et al., 2002). Here, we document a $\sim 3$‰ offset in authigenic clay ${\delta }^{18}O$ between the western and eastern facies, with lower values in the east. This offset is best explained by the influence of the Blue Mountains on local precipitation patterns, indicating the range has supported local precipitation gradients since at least the Eocene.

Geologic Setting

Prior to the deposition of Cenozoic sediment in the John Day region, the Mesozoic accretion of the Wallowa and Olds Ferry arcs created the Blue Mountains province (Dickinson and Thayer, 1978; Dickinson, 1979; Schwartz et al., 2010; LaMaskin et al., 2015). The province accreted along the Salmon River suture zone, likely sometime before 130 Ma (Dickinson, 1979, Dickinson, 2004; Schwartz et al., 2010; LaMaskin et al., 2015). After accretion, deformation of the suture zone is thought to have ended around 90 Ma, setting the stage for Cenozoic deposition within the John Day region.

Deposition in the John Day region occurred throughout the Eocene to the Pliocene and sediment was predominantly derived from air-fall ash and some ash-flow sheets with evidence for minor reworking by alluvial and lacustrine processes (Swanson and Robinson, 1968; Robinson et al., 1984; Graham, 2014). Volcanism appears frequent throughout the depositional record and likely shifted westward from the Blue Mountains province to the proto-Cascades ∼37–40 million years ago, possibly due to “flat slab” detachment and a steepening of western North American subduction (Lipman et al., 1972; Noble, 1972; Robinson et al., 1984; Heller et al., 1987). This shift may be linked to the start of Western Cascades uplift, but early volcanism in the Western Cascades was likely associated with isolated stratovolcanoes along a low-lying coastal plain that may not have intercepted westerly moisture as effectively as today’s Cascades Range (White and Robinson, 1992; Dillhoff et al., 2009). The timing of the onset of the Cascades rainshadow is still debated, and likely varies north-south (Priest, 1990; Kohn et al., 2002; Takeuchi and Larson, 2005; Takeuchi et al., 2010; Bershaw et al., 2020; Pesek et al., 2020), but drier, rainshadow-like conditions appear to emerge in Oregon in the late Oligocene to early Miocene (Woodburne and Robinson, 1977; MacFadden and Hulbert, 1988; Retallack, 2004).

The eastern and western facies of John Day deposition are separated by the Blue Mountains but they share some lithologic similarities (Swanson and Robinson, 1968; Robinson et al., 1984). The facies share the same formations for most of the Cenozoic, from the lower-mid Eocene Clarno Formation to the upper Eocene–lower Miocene John Day Formation. After deposition of the John Day Formation, the western facies contains the Ellensburg, Simtustus, Dalles, and Deschutes Formations while the eastern facies contains the Mascall and Rattlesnake Formations (Smiley, 1963; Farooqui et al., 1981; Smith, 1986; Graham, 2014). Both facies are primarily comprised of claystones and air-fall tuffs. The western facies contains more ash-flows and lava-flows than the eastern facies, suggesting a volcanic source west of the Blue Mountains and indicating the topographic barrier of the Blue Mountains may have limited the eastern extent of these deposits (Swanson and Robinson, 1968; Robinson et al., 1984). A third unit, referred to as the southern facies, also exists within the Blue Mountains region and shares lithological similarities with the eastern facies, but was not sampled in this study. The western facies, especially after the westward shift in volcanism at ∼40 Ma, generally hosts coarser grained volcaniclastic material with thicker sedimentary packages, further supporting a western source of volcanic material (Robinson et al., 1984).

Paleobotanical evidence in the John Day region points to a long-term aridification trend beginning about 30 million years ago with the expansion of open-habitat grasslands. Paleosols at this time contain evidence for grassy root textures, and mammal fossils reveal adaptations for running (“cursoriality,” a common indicator of open vegetation) and eating tougher foods like grasses (“hypsodonty”) (MacFadden and Hulbert, 1988; Janis et al., 2002; Retallack, 2004). This transition to drier conditions, lasting until ∼19 Ma, is generally interpreted to reflect strengthening of the Cascades rainshadow and/or a shift in the seasonality of precipitation from summer-dominated to the winter-dominated regime that exists today (Retallack, 2004; Retallack, 1997).

Methods

Authigenic Clay Sample Preparation and Isotopic Analysis

We analyzed the oxygen isotope composition of 29 samples spanning 15 localities to build on existing data from Takeuchi (2007). Clay-rich paleosol samples were gently ground with a mortar and pestle and centrifuged to separate the <0.5 µm size fraction. The separated material was dried and gently ground with a mortar and pestle into a powder. Samples with indication of organic matter (dark coloration) were treated with a 3% hydrogen peroxide (H₂O₂) solution for 24–48 h before being rinsed at least 5 times with de-ionized water in a centrifuge.

At least one sample from each sedimentary package at each locality was run for X-ray diffraction to identify the clay mineralogy and screen for quartz. X-ray diffractometry was run at the Stanford University Environmental Measurements Facility using a Rigaku MiniFlex 600 Benchtop X-ray Diffraction System. The diffractometer is equipped with a copper (Cu) anode set at the maximum power of 600 W. Powdered samples were first suspended in isopropanol and left to air-dry on a zero-background quartz sample holder. Once dried, each sample was run twice with the 2-theta angle ranging from 2 to 90°. Samples were untreated in the first run and glycolated in the second run. To glycolate the samples, dried powders were placed in a sealed desiccator with 1 cm of ethylene glycol at the base and left overnight in an oven set to 65°C. The Rigaku PDXL software was used to aid in mineral identification for each sample. Samples identified to have quartz and/or illite peaks were not analyzed for oxygen isotopes.

After screening for quartz and illite, 1–2 mg of sample powder was mixed with lithium flouride ($LiF$; $\sim$1:1 by mass) and pressed into pellets. The $LiF$ pellets prevent dispersion of clay powder during laser ablation. Samples were brought to a vacuum to remove atmospheric vapor in the $LiF$. Any remaining vapor and any sorbed water in smectite minerals was removed through 2–4 pre-flourinations where samples were exposed to bromine pentaflouride ($Br{F}_5$) for ∼90 s. Afterward, samples were laser ablated with a New Wave Research MIR10-25 infrared laser ablation system in the presence of $Br{F}_5$ to liberate $O_2$ gas (Sharp, 1990; Sjostrom et al., 2006; Mix and Chamberlain, 2014; Mix et al., 2016). Oxygen isotope analyses on $O_2$ gas were performed in dual-inlet mode at the Stanford University Stable Isotope Biogeochemistry Laboratory on a Thermo Finnigan MAT 252 or 253+, depending on the date of acquisition (Abruzzese et al., 2005; Hren et al., 2009; Mix et al., 2016; Chamberlain et al., 2020). We purified $O_2$ gas from samples run on the MAT 252 with two liquid nitrogen cold traps and one potassium bromide ($KBr$) trap before freezing $O_2$ on a liquid nitrogen-temperature zeolite and subsequently equilibrating it with the mass spectrometer sample bellows. On the MAT 253+, $O_2$ gas was purified with three liquid nitrogen traps and one sodium chloride ($NaCl$) trap, frozen on a zeolite at liquid nitrogen temperature, passed over a room temperature flow-through zeolite with high purity He as the carrier gas, and finally frozen in a zeolite trap and equilibrated with the mass spectrometer sample bellows. Repeated analyses of in-house standard DS069 were made during each day of analysis to correct for drift. Precision of DS069 replicates was 0.26‰ (n = 43). All isotopic ratios are reported relative to Vienna Standard Mean Ocean Water (VSMOW).

Comparing Oxygen Isotope Values of the Eastern and Western Facies

In order to map out the spatial pattern of authigenic clay ${\delta }^{18}O$ through the Cenozoic we calculate the residual of all data points relative to the Cenozoic trend. We build this Cenozoic trend by taking the average of the eastern and western facies trends, each defined by a loess smooth curve. This approach ensures that the Cenozoic trend is not biased toward one facies or the other when its sample density is relatively higher for a given point in time. After establishing the Cenozoic trend we calculate the residual ${\delta }^{18}O$ value for each sample (${\delta }^{18}{O}_{sample}-{\delta }^{18}{O}_{trend}$). This represents a detrended record of ${\delta }^{18}O$ that allows for a spatial comparison of samples from different times of the record.

Results

Most clay-size separates were comprised of smectite or mixed smectite/kaolinite with few kaolinite samples. Samples containing illite and/or quartz were not analyzed for ${\delta }^{18}O$. Kaolinite-rich samples are mostly found in the Eocene when conditions were likely wetter, consistent with the findings of Takeuchi (2007). Typical diffraction patterns for smectite, kaolinite, and excluded (illite/quartz) samples can be found in Supplementary Figure S1. We plot oxygen isotope data of kaolinite and smectite samples together because their fractionation factors are similar for at environmental temperatures (Sheppard and Gilg (1996); Supplementary Figure S3). We note that a recent database for oxygen isotope fractionation suggests a 2–3‰ offset between kaolinite and smectite fractionation at environmental temperatures (Vho et al., 2019), but the database is not recommended for use at low, environmental temperatures because a key approximation does not hold in this range Eq. 5 of Vho et al. (2019)).

Oxygen isotope values of our composite record range from 8.3 to 16.9‰ with an average value of 12.5‰ (±2.2‰, 1 σ). In the eastern facies the average ${\delta }^{18}O$ value is 11.1‰ (±1.9‰, 1 σ) and in the western facies mean ${\delta }^{18}O$ is 13.7‰ (±1.6‰, 1 σ). Throughout the record, oxygen isotope values in the eastern facies (the Blue Mountains province) are ∼3‰ lower than those from the western facies, while ${\delta }^{18}O$ values in both facies follow similar trends. Specifically, eastern and western facies ${\delta }^{18}O$ values both show a ∼3‰ increase between 30 and 20 Ma (Figure 2A).

FIGURE 2. (A) Oxygen isotope values of authigenic clay samples (‰ VSMOW) in the eastern (dark blue) and western (light blue) facies. Diamonds denote data from this study and circles are from Takeuchi (2007). In both panels the thin, horizontal and vertical gray lines are age uncertainty and δ¹⁸O uncertainty, respectively. Thick, gray line denotes the average of the eastern and western facies trends, representing a regional trend. (B) Residual δ¹⁸O values relative to the regional trend (gray line). Eastern facies δ¹⁸O values are ∼3‰ lower than the western facies throughout the record.

Discussion

Oxygen Isotopes of the Eastern and Western Facies

Lower oxygen isotope values in the Blue Mountains province (eastern facies) suggests this local topographic feature has influenced regional precipitation patterns for at least the last ∼50 million years (Dickinson and Thayer, 1978; Dickinson, 1979; Schwartz et al., 2010; LaMaskin et al., 2015). Based on spatial interpolation of oxygen isotope residuals (Figure 2B) we find that the boundary between negative (lower ${\delta }^{18}O$) and positive (higher ${\delta }^{18}O$) residuals closely tracks the modern boundary of the eastern and western facies (which is, itself, defined by the extent of the Blue Mountains) (Figure 3). This close correspondence with the modern extent of the Blue Mountains suggests the possibility that authigenic clay ${\delta }^{18}O$ values reflect modern conditions due to diagenetic alteration or overprinting by modern waters instead of reflecting changes in past precipitation ${\delta }^{18}O$. However, ${\delta }^{18}O$ values from the eastern and western facies both increase by ∼3‰ between 30 and 20 Ma, and this shift likely would not be coherent if all samples have been similarly altered by modern waters. Thus, we interpret ${\delta }^{18}O$ values to represent the oxygen isotope composition of water at the time of clay mineral formation.

FIGURE 3. (A) Zoomed-in panel (see Figures 1B,C) of sample locations (black diamonds) and mean annual precipitation. Light and dark blue polygon outlines reflect the approximate spatial extent of the western and eastern facies, respectively (after Albright et al. (2008); Graham (2014); Robinson et al. (1984); Takeuchi (2007)). (B) Spatially interpolated map of ${\delta }^{18}O$ residuals overlain with eastern and western facies extents.

Today, the windward slopes of the Blue Mountains capture more annual precipitation than the adjacent high plains (see Figure 1C) and receive a larger fraction of annual precipitation during the winter months (Supplementary Figure S2). Both of these observations support lower ${\delta }^{18}O$ values in the Blue Mountains than in the adjacent plains. First, increased rainout over mountains is often explained by topography forcing air upward, leading to adiabatic cooling and moisture condensation and precipitation (Aristotle, 1931; Smith, 1979; Smith and Barstad, 2004; Roe, 2005). Higher precipitation due to Blue Mountains orography implies more rainout and a decrease in ${\delta }^{18}O$ values as precipitation preferentially removes ${\delta }^{18}O$. Alternatively, lower ${\delta }^{18}O$ values in the Blue Mountains region may be related to precipitation seasonality as, compared to the plains, the Blue Mountains receive a greater fraction of total precipitation in the winter months when precipitation ${\delta }^{18}O$ values are lower. It is possible that upslope rainout and precipitation seasonality both contribute to lower ${\delta }^{18}O$ values in the Blue Mountains. However, disentangling the effect of seasonality from upslope rainout on precipitation ${\delta }^{18}O$ is difficult because most modern ${\delta }^{18}O$ data in the John Day region are sampled in warmer months or from rivers that integrate seasonal precipitation (e.g., Bershaw et al. (2019)).

Despite the close link between residual ${\delta }^{18}O$ values and the extent of the Blue Mountains, ${\delta }^{18}O$ residuals do not correlate strongly with modern elevation or fraction of winter precipitation (Supplementary Figure S4). Residual ${\delta }^{18}O$ values generally decrease with elevation by ∼3.3‰ per kilometer, similar to the regional lapse rate of ∼3.2‰ Bershaw et al. (2020) per kilometer, but the correlation is weak ($R^2=0.17$). There is no significant correlation between residual ${\delta }^{18}O$ and the fraction of annual precipitation occurring in winter. The weak relationships between past clay ${\delta }^{18}O$ and modern conditions is not unexpected since some clays likely formed in contact with local meteoric water (tracking local elevation and precipitation seasonality) while others formed in floodplains in contact with upstream runoff (tracking the upstream hypsometric mean elevation and precipitation seasonality). Without reliable constraints on past floodplain extents and local drainage divides it is difficult to determine the relative importance of elevation, precipitation seasonality, or other factors in eastern and western facies ${\delta }^{18}O$ values. Additionally, modern winter precipitation and topography may differ from that of most samples in our record. Thermochronometry and geochronology data from the western Cascades in the latitude range of our samples points to exhumation between 20 and 10 Ma (Pesek et al., 2020), and there is abundant evidence that the timing of Cascades uplift and exhumation varied north-to-south (Reiners et al., 2002; Takeuchi and Larson, 2005; Pesek et al., 2020). If spatially variable topographic changes affected the spatial pattern of ${\delta }^{18}O$, a direct comparison of past John Day data with modern conditions may be invalid.

Implications for Regional Climate and Tectonics

Our results suggest that a regional gradient in precipitation amount and/or seasonality has existed around the Blue Mountains for most of the last 50 million years. If the Blue Mountains did not increase rainout or were not elevated above the plains, we would expect no difference in clay ${\delta }^{18}O$ values between the eastern and western facies. While the Blue Mountains are generally wetter than the adjacent plains today, lower ${\delta }^{18}O$ values in the eastern facies does not require that the Blue Mountains were wetter in the past. For example, it is possible that the Blue Mountains and the plains received similar amounts of annual precipitation, but summer (winter) precipitation made up a larger fraction of annual precipitation in the plains (mountains). However, summer and winter precipitation both increase with elevation in the Blue Mountains (and across the western United States; Supplementary Figure S5) today, and there is no clear reason why topography would not have the same effect on precipitation in the past. We suggest the Blue Mountains have received more precipitation than the surrounding plains for at least the last ∼50 million years. Thus, a possible Eocene boundary between tropical lowland and temperate highland vegetation (Bestland and Retallack, 1994; Bestland et al., 2002) is likely driven by colder temperatures or topographic relief, but not drier conditions in the Blue Mountains.

Despite a long-lived precipitation gradient associated with the Blue Mountains, both the Blue Mountains and the plains record an increase in ${\delta }^{18}O$ values that tracks independent evidence for regional aridification. This ${\delta }^{18}O$ increase is probably not driven by temperature since the shift is unidirectional while global climate both warmed and cooled between 30 and 20 Ma (Zachos et al., 2001). Meanwhile, the onset of drier conditions between 30 and 20 million years ago is evidenced by the expansion of open-habitat grasslands and mammals adapted to running and eating tougher vegetation like grasses (Woodburne and Robinson, 1977; MacFadden and Hulbert, 1988; Jacobs et al., 1999; Janis et al., 2002; Retallack, 2004), although this transition may also be explained by grassland-grazer coevolution, independent of climate (Retallack, 2001; Retallack, 2013). The ∼3‰ increase in authigenic clay ${\delta }^{18}O$ may be related to drying, but the interpretation is not straightforward. For example, Retallack (2004) suggests that drying is related to 1) the uplift of the Cascades and 2) a decrease in summer (high-${\delta }^{18}O$) precipitation. But both of these effects would likely decrease ${\delta }^{18}O$ values rather than increase them. Further, marine ${\delta }^{18}O$ values vary by just 1‰ from 30 to 20 Ma and do not show a systematic increase capable of overprinting the effects of Cascades uplift and less summer precipitation (Zachos et al., 2001). Increased upstream (westward) moisture recycling in drier climates can also increase ${\delta }^{18}O$ (Salati et al., 1979; Sonntag et al., 1983; Mix et al., 2013; Chamberlain et al., 2014; Winnick et al., 2014; Kukla et al., 2019), but this effect is negligible in near-coastal settings where upstream land area is minimal (Mix et al., 2013; Winnick et al., 2014; Kukla et al., 2019).

Instead, we speculate that the ${\delta }^{18}O$ increase in our data reflects the onset of the Cascades rainshadow driving a decrease in winter (low-${\delta }^{18}O$) precipitation. Today, the Cascades represent a much stronger rainshadow in the winter than in the summer (e.g., Siler and Durran (2016); Supplementary Figure S6), suggesting the possibility that their uplift would disproportionately decrease winter precipitation inland. Still, it is not clear that the increase in ${\delta }^{18}O$ from less winter precipitation could outweigh any decrease in ${\delta }^{18}O$ from Cascades uplift. It is possible that other factors associated with drier conditions, like surface water and subcloud evaporation, could also increase ${\delta }^{18}O$. Thus, more work is needed to test the hypothesis that drying in the John Day region is owed to disproportionate drying in the winter.

The distinct isotopic signature of the Blue Mountains emphasizes the importance of local elevation in regional tectonic reconstructions. Even though the eastern and western facies lie leeward (east) of the Cascades, lower ${\delta }^{18}O$ values in the eastern facies (the Blue Mountains) may be misinterpreted to reflect the height of the Cascades instead of local topography. For example, using hydrogen isotopes of a volcanic glass sample from the Blue Mountains, Bershaw et al. (2019) suggests the Cascades may have been higher than present as early as ∼32 million years ago. Alternatively, we suggest that low hydrogen isotope ratios in the Oligocene reflect the local topography of the Blue Mountains rather than upstream elevation of the Cascades. Interestingly, the hydrogen isotope data of Bershaw et al. (2019) reveal an increase in $\delta D$ of ∼16‰ (∼2‰ in ${\delta }^{18}O$) that happens near the mid-Miocene, between ∼16 and 8 Ma, significantly later than the increase of ∼3‰ between 30 and 20 Ma found in our data. Due to sparse hydrogen isotope data east of the Cascades older than 20 Ma, it is not clear whether volcanic glass $\delta D$ also records an increase from 30 to 20 Ma. Further, while there are few coeval glass samples from the eastern and western facies to test for a spatial pattern, the existing data do not show a systematic difference in $\delta D$ between the two regions. If the spatial and temporal trends of authigenic clay and volcanic glass isotopes differ, it is likely that the proxies are forming under different conditions and may be seasonally biased relative to one another. Such information would provide increasingly nuanced insight to the uplift history of the Cascades and its effect on leeward precipitation seasonality. However, at present there is not enough overlap between our authigenic clay data and the volcanic glass data to determine whether the results truly reflect distinct trends.

Conclusion

The Blue Mountains have likely supported local (<100 km) gradients in precipitation for much of the Cenozoic with more precipitation reaching the Blue Mountains than the surrounding plains. Distinct floral assemblages from the same stratigraphic interval provided the first hint that the John Day region may have hosted multiple biomes as early as the Eocene (Bestland and Retallack, 1994; Bestland et al., 2002), and our stable isotope results suggest these biomes could co-exist due to distinct elevation and climate regimes. We also find no evidence that regional drying from 30 to 20 million years ago was driven by a shift from summer-dominated to winter-dominated precipitation (Retallack, 2004).

Our high spatial resolution constraints on Blue Mountains topography provides a benchmark for future paleo -floral and -faunal work in the John Day region. Efforts to compare floral and faunal assemblages from the eastern and western facies will constrain how the Blue Mountains influenced local ecosystem dynamics through time. Additionally, our results suggest that the comparison of authigenic clay ${\delta }^{18}O$ and mammal tooth enamel ${\delta }^{18}O$ may allow for constraints on local mammal migration patterns. For example, comparing tooth enamel ${\delta }^{18}O$ from mammals of the same site may inform whether the range of some animals was more local than others (living in the eastern facies, western facies, or traveling between both). Further, changes in the inferred seasonal amplitude of tooth enamel ${\delta }^{18}O$ could reflect changes in seasonal migration between the two facies. Overall, our constraints on the spatial pattern of Blue Mountains topography and its possible influence over regional precipitation opens new opportunities for biogeography research on the relationship between landscapes and life in John Day through the Cenozoic.

Data Availability Statement

The original contributions presented in the study are included in the article/Supplementary Material, further inquiries can be directed to the corresponding author.

Author Contributions

TK, DI, JR, SK, and CC conceptualized the study. TK, DI, and CM performed laboratory analysis. TK, DI, JR, JG, DM, and CC performed fieldwork. JG and DM led field-based observation. TK analyzed data and wrote the original draft. All authors contributed to editing and revising.

Funding

This research was funded by NSF EAR-1322084 and Heising-Simons grants to CPC.

Conflict of Interest

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Acknowledgments

The authors acknowledge J. Bershaw and Gregory Retallack for comments that improved the manuscript. We also acknowledge Joshua X. Samuels, Nicholas Famoso, Shelley Hall, and Patrick Gamman for assistance obtaining research permits at John Day Fossil Beds National Monument. We also thank Joshua X. Samuels for assistance with field stratigraphy and sampling, and Nicholas Famoso for further guidance on sampling and our age model. Elevation data are sourced from Terrain Tiles hosted on Amazon Web Services. Terrain Tiles data combines Global ETOPO1 and 3DEP datasets. Global ETOPO1 terrain data is courtesy of the U.S. National Oceanic and Atmospheric Administration. United States 3DEP (formerly NED) and global GMTED2010 and SRTM terrain data are courtesy of the U.S. Geological Survey. Research at John Day Fossil Beds National Monument was conducted under Study Number JODA-00033 with Permit Numbers JODA-2017-SCI-001 and JODA-2018-SCI-002, issued by John Day Fossil Beds and the United States Department of the Interior National Park Service.

Supplementary Material

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/feart.2021.635181/full#supplementary-material.

References