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PaleoBios 28(3):89–113, January 22, 2009© 2009 University of California Museum of Paleontology
Cenozoic cooling and grassland expansion in Oregon and Washington
GREGORY J. RETALLACKDepartment of Geological Sciences, University of Oregon, Eugene, OR 97403; [email protected]
Many different kinds of paleosols are common in Cenozoic badlands of the high deserts of eastern Oregon and Washington. Pedotypes, taxonomic orders, and climofunctions are three distinct approaches that use paleosol data to reconstruct past climate, ecosystems, topography, parent materials, and landscape stability. Because the back-arc paleotopographic setting, rhyolitic parent material, and rate of subsidence changed little for paleosols of eastern Oregon and Washington over the past 45 million years, these paleosols record primarily changing climate and vegeta-tion. Paleosols of Oregon and Washington are evidence of the transition from a late Eocene (35 Ma) peak of warm, humid forests to cool, desert shrublands in the Oligocene. The reversion to warm-wet forests in the middle Miocene (16 Ma) was followed by general cooling and spread of sod grasslands that culminated in the last ice age (>1.8 Ma). These records are compatible with a greenhouse model for climate change of cooling induced by coevolutionary advances in carbon sequestration and consumption by grassland soils and sediments, punctuated by warming from volcano-tectonic and extraterrestrial perturbations. The abundant Cenozoic paleosols in the Pacific Northwest are multifaceted guides to past global change.
INTRODUCTION
Oregon’s Painted Hills of red-streaked badlands are the best known of colorful badlands scattered throughout the high desert of eastern Oregon and Washington (Figs. 1, 2B). In the Painted Hills, the prominent red bands are clayey subsurface (Bt) horizons of forest soils (Alfisols) representing successive humid-climate ecosystems that colonized tuffa-ceous alluvial terraces over periods of tens of thousands of years (Retallack et al. 2000). Evidence for this interpretation comes from studies of root traces (drab-haloed claystone),
soil structures (angular blocky peds and sesqui-argillans), soil microfabrics (porphyroskelic skelmasepic), grain-size distri-bution (subsurface clay enrichment) and chemical composi-tion (alkali depletion and iron oxidation). The Painted Hills includes at least ten paleosols of this Luca pedotype, as well as 14 other kinds of paleosols that are among 435 paleosols in a composite stratigraphic section (430 m thick) of the lower John Day Formation. The abundance and diversity of paleosols in the Painted Hills is not unusual (Retallack et al. 2002a, Bestland and Forbes 2003, Sheldon 2003, Retallack 2004a, b, Retallack et al. 2004, Sheldon 2006), and calls for a systematic approach to classifying, naming, and interpret-ing the thousands of paleosols known from the Cenozoic of the Pacific Northwest. This paper reviews the long record of paleosols in Oregon and Washington, comparing different approaches of: (1) their field classification (2) their interpreta-tion within modern soil taxonomies; and (3) the application of transfer functions derived from study of modern soils. Such varied approaches to paleosol archives have implications for understanding long-term changes in climate and vegeta-tion, both locally and globally (Retallack 2007). This paper emphasizes the paleosol record and different approaches to its interpretation.
COMMON PALEOSOLS OF THE PACIFIC NORTHWEST
Cenozoic fossils of Oregon have been famous since Thomas Condon’s discovery in 1865 of fossil leaves and mammal bones in what is now known as the Painted Hills Unit of the John Day Fossil Beds National Monument (Clark 1989). University of California paleontologists John Merriam (1906, 1916, Merriam and Sinclair 1907), Ralph Chaney and Dan Axelrod (Chaney, 1924, 1948, Chaney and Axelrod 1959) subsequently collected rich fossil floras and faunas from Eocene, Oligocene, and Miocene rocks nearby, and geologists Hay (1962) and Fisher (1964) were the first
Figure 1. Localities of paleosols examined in Oregon and Wash-ington.
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to note paleosols in these rocks. The temporal resolution and completeness of this paleosol record recently became apparent when the geochronology of these strata was refined by magnetostratigraphy (Prothero and Rensberger 1985, Prothero et al. 2006a, b) and by radiometric dating of nu-merous interbedded volcanic tuffs and lava flows (Fremd et al. 1994, Streck et al. 1999). With these new data has come an appreciation of the abundance and diversity of paleosols, outlined below.
Clarno Formation
The Hancock Field Station near Clarno has a sequence of at least 76 successive paleosols of 11 different types de-veloped within volcaniclastic conglomerates and tuffaceous red beds of the Clarno Formation that accumulated on the margins of a large andesitic stratovolcano (Retallack et al. 2000). The Nut Beds Member of the Clarno Formation has yielded a middle Eocene flora of 173 species of fossil fruits and seeds from tropical forest plants, including magnolia, palms, and vines (Manchester 1994). Fossils from near Clarno also include permineralized wood (Wheeler and Manchester 2002) and middle Eocene (Bridgerian) mammals, as well as a later Eocene (Duchesnean) bone bed yielding titanotheres, creodonts, and other mammals (Hanson 1996). Radiometric dating of andesitic and basaltic flows and rhyodacitic tuffs at the Clarno Nut Beds locality allows age determination by linear interpolation from the following tie points: 43.8 ± 0.5 Ma basalt at 38 m, 43.0 ± 3.5 Ma tuff at 52 m, 42.7 ± 0.3 Ma tuff at 90 m, 39.2 ± 0.03 Ma tuff at 128 m (Bestland et al. 1999, Retallack et al. 2000).
John Day Formation
The Painted Hills Unit of the John Day Fossil Beds Na-tional Monument exposes the lower John Day Formation, with hundreds of paleosols ranging from deeply weathered, red claystones, to little-weathered volcanic siltstones (Figs. 1A–B). Only calcareous volcanic soils and associated sedi-ments have yielded Oligocene (early Arikareean) mammals (Retallack et al. 2000). Bridge Creek near the Painted Hills has lent its name to fossil floras of temperate paleoclimatic affinities, dominated by dawn redwood, alder and oak (Meyer and Manchester 1997). High quality 40Ar/39Ar single-crystal laser radiometric dates of tuffs in the Painted Hills range from 40–29 Ma (late Eocene to middle Oligocene): 39.7 ± 0.03 Ma at 0 m, 32.99 ± 0.11 Ma at 98 m, 32.66 ± 0.03 Ma at 137 m, 29.75 ± 0.02 Ma at 329 m, 29.70 ± 0.06 Ma at 381 m, and 28.65 ± 0.05 Ma at 427 m (Retallack et al. 2000). There are hundreds more unstudied paleosols in the upper John Day Formation in this area, which continues from Bridge Creek north into the canyons south of Sutton Mountain.
Paleosols and mammal fossils of the upper John Day Formation are best known from the Sheep Rock and Blue Basin units of the John Day Fossil Beds north to Longview Ranch, Johnson Canyon, Kimberly and Spray (Retallack 2004a). These paleosols contain calcareous nodules and
preserve abundant mammal bones and teeth (Retallack 1998). Orellan(?), Arikareean, and Hemingfordian mammals (Coombs et al. 2001, Hunt and Stepleton 2004), radiometric dates of tuffs (Fremd et al. 1994), and magnetostratigraphy (Prothero and Rensberger 1985) indicate that this sequence ranges in age from 29–17 Ma (middle Oligocene to early Miocene). Beginning at top of tuff H on Longview Ranch, recent 40Ar/39Ar single-crystal laser dates include 28.7 ± 0.07 Ma at 0 m, 27.89 ± 0.57 Ma at 67 m, 27.18 ± 9,13 Ma at 106 m, 25.9 ± 0.31 Ma at 177.5 m, and 19.6 ± 0.13 Ma at 497 m (Fremd et al. 1994, Retallack 2004a, Tedford et al., 2004).
Columbia River Basalt
The Middle Miocene (16 Ma) Columbia River Basalt Group includes numerous paleosols between the flood ba-salts. These have been studied extensively in Picture Gorge (Sheldon 2003, 2006). Although these red paleosols have been regarded as baked zones from the heat of overlying flows, the vitrinite reflectance of peaty interbasaltic paleosols and the degree of ceramicization of red clayey inter basaltic paleosols both indicate that only the top few centimeters of the meter-thick paleosols have been thermally altered (Sheldon 2003). Vaporization of vegetation thoroughly vesiculated and zeolitized the flows, rather than altering underlying paleosols. Tie points for ages within the section beginning 300 m north of Picture Gorge include a 16.01 Ma magnetic reversal at 50 m, and the 15.77 ± 0.04 Ma Mascall tuff at 280 m (Watkins and Baksi 1974, Bailey, 1989, Tedford et al., 2004).
Mascall Formation
The middle Miocene Mascall Formation near Dayville is well known for a fossil flora of oak and bald cypress within diatomaceous lake beds (Chaney and Axelrod 1959). Paleo-sol-rich overbank deposits of the Mascall Formation yield Barstovian mammal assemblages in the type section near Day-ville, through Rock Creek and Antone, and west to Paulina (Downs 1956). Bones and teeth are rare because paleosols of the Mascall Formation are clayey and non-calcareous, more like those of the lower than upper John Day Forma-tion (Bestland and Forbes 2003). Chronostratigraphic ages of paleosols for the lower part of the formation in the type area include a 16.2 ± 1.4 Ma ash at 35 m, 15.77 ± 0.04 Ma ash at 119 m, 15.2 Ma magnetic reversal at 120 m, 15.0 Ma reversal at 129 m, 14.9 Ma reversal at 165 m, and 14.8 Ma reversal at 190 m (Tedford et al., 2004, Prothero et al. 2006a, Prothero et al. 2006b). Other Barstovian (middle Miocene) paleosols are known from fossil mammal sites near Gateway (Downs 1956, Ashwill 1983, Smith 1986) and Beatty Buttes (Wallace 1946, Hutchison 1968).
Ironside and Juntura Formations
Late Miocene plant, bird, and mammal fossils of the Clarendonian land mammal age have been reported from the Unity, Ironside, and Juntura formations (Bowen et al. 1963, Shotwell 1963, Shotwell and Russell 1963, Retallack,
RETALLACK—CENOZOIC COOLING AND GRASSLAND EXPANSION IN OR AND WA 91
2004b). A sequence of tuffs in the Ironside Formation of upper Windlass Gulch, north of Unity have been correlated with tuffs radiometrically dated as 11.8 Ma at 49.5 m, 11.8 Ma at 60.0 m, 11.7 Ma at 73.6 m, 11.7 Ma at 83.3 m, 11.6 Ma at 100.5 m, 11.6 Ma at 110 m, 11.5 Ma at 115.5 m, and 11.3 Ma at 173 m (Perkins et al. 1998, Retallack 2004b). Radiometrically dated rocks in the Juntura Formation west of Juntura include a 12.4 ± 0.5 Ma basalt at 0 m, 11.5 ± 0.6 Ma tuff at 230 m, and 9.7 ± 0.22 Ma ash-flow tuff at 259 m (Bowen et al. 1963, Evernden and James 1964, Fiebelkorn et al. 1982, Johnson et al. 1996, Retallack 2004b, Tedford et al., 2004). Unity, Ironside, and Juntura paleosols are weakly calcareous and siliceous like soils found in the summer dry grassy woodlands of northern California today (Retallack 2004b).
Rattlesnake, Deschutes, Shutler, and Drewsey Formations
The latest Miocene Rattlesnake Formation near Dayville yields Hemphillian mammals, including large grazing horses (Merriam et al. 1925), and chronostratigraphic ages that include a 7.14 Ma magnetic reversal at 25 m, a 7.05 ± 0.01 Ma tuff at 77 m, and a 7.0 Ma reversal at 101 m (Streck et al. 1999, Prothero et al. 2006b). Paleosols of the Rattlesnake Formation reveal a transition from woodland through tall grassland to sagebrush steppe (Retallack et al. 2002a). The Rattlesnake Ash-Flow Tuff is widespread beyond the type area near Dayville, and overlies paleosols at Logan Butte and in Spanish Gulch (Retallack 2007). Paleosol sequences of the Deschutes Formation near Prineville have the 7.05 ± 0.01 Ma ash-flow tuff at 0 m, and a 3.3 ± 0.2 Ma basalt rimrock at 103 m (Bishop and Smith 1990, Streck et al. 1999).
Other Hemphillian mammals found near Rome are corre-lated with the section at Juntura, for which radiometric dates are 9.7 ± 0.22 Ma at the base of the Drewsey Formation and 8.08 ± 0.56 Ma for a basalt at 100 m (Wolf and Ellison 1971, Fiebelkorn et al. 1982, Sheppard and Gude 1987). The Shut-ler Formation at McKay Reservoir yields a late Hemphillian mammal assemblage (Shotwell 1956, 1958).
Ringold Formation
The Miocene-Pliocene Ringold Formation yields fragmen-tary Blancan fossil mammals, and is well exposed in White Cliffs along the Columbia River near Richland, and in road and railway cuts near Taunton, Washington (Gustafson 1978, Smith et al. 2000). The White Cliffs section records the fol-lowing paleomagnetic reversals: 5.23 Ma at 2.81 m, 4.98 Ma at 32 m, 4.89 Ma at 40 m, 4.8 Ma at 44 m, 4.62 Ma at 63.1 m, 4.48 Ma at 61.1 m, and 4.29 Ma at 142 m (Gustafson 1985). Near Taunton, the formation records the Pliocene paleomagnetic reversal at 3.040–3.110 Ma between 0–9.9 m in a local section above a railway siding (Morgan and Morgan 1995, Smith et al. 2000). Plio-Pleistocene paleosols are pres-ent in the 2-Ma Huckleberry Tuff exposed on high terraces south of Sutton Mountain (Perkins et al. 1998). Paleosols at
White Cliffs, Taunton, and Sutton Mountain include crumb-structured soils like those of grasslands, as well as calcareous nodular soils like those of sagebrush steppe.
Palouse Loess
The Palouse Loess is an eolian deposit of redeposited rhyodacitic ash and quartzofeldspathic fluvial silt . It forms a thick and agriculturally fertile veneer on the Columbia River Basalts throughout much of the Columbia River basin in eastern Oregon and Washington. Dating by magnetostratig-raphy, tephrochronology, and thermoluminescence of the road cut west of Helix, Oregon, includes ages of 0.0636 Ma at 4 m and 0.158 Ma at 6 m from the top (Tate 1998). An outcrop near Washtucna, Washington yielded a paleomagnetic age of 0.788 Ma at 12 m and a radiometric age 0.055 Ma for an ash layer at 3 m from the top (Busacca 1989, 1991). The best known sections near Helix and Washtucna show Milankovich cyclicity (41–100 ka) alternating between two distinct pedotypes: (1) crumb-textured paleosols with abun-dant earthworm microfabrics like those in grassland soils, and (2) massive calcareous paleosols riddled with cicada burrows like those of sagebrush steppe soils (Tate 1998, O’Geen and Busacca 2002, Blinnikov et al. 2002).
CLASSIFICATION OF PALEOSOLS
Paleosols can be classified by field pedotypes or taxonomic orders. Pedotypes are different kinds of paleosols based on specified type sections, and recognizable from field charac-teristics. In contrast, taxonomic orders are categories of clas-sification based on modern soils, which can only be applied to paleosols after consideration of laboratory data obtained as proxies for taxonomic criteria (Retallack 1997a).
Pedotypes
The general paleosol type, or pedotype, is a non-genetic field mapping designation (Retallack 1994), comparable with “series” of soil maps (Soil Survey Staff 1993). Each pedotype is based on a specified type profile (Retallack et al. 2000, 2002a, Retallack 2004a, b). Cenozoic pedotypes in Oregon and Washington have been named using simple descriptive terms from the Sahaptin Native American language (Figs. 2–3), which was widely spoken throughout the southern Co-lumbia River basin (Rigsby 1965, De Lancey et al. 1988).
The purpose of pedotype naming and description is to facilitate precise description and characterization of paleo-sols as natural objects in their own right. Conventions of naming and definition, such as nomination of type sections, are needed to facilitate the growth of descriptive databases through the efforts of successive investigators. These conven-tions follow standard procedures of soil survey (Soil Survey Staff 1993). Differences between pedotypes need to be documented (Fig. 3) before their environmental significance is interpreted. Although interpretations of paleoclimate, fossil flora and fauna, paleotopographic position, parent material, and time for formation differ between pedotypes, they can
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be generalized to different individual paleosols of the same pedotype (Retallack et al. 2000).
Pedotype interpretation
Each pedotype represents a particular combination of soil-forming influences; i.e., a particular ecosystem formed in a particular climate, on a particular parent material, in a particular landscape setting, and over a particular period of time. As with paleontological data, pedotypes have lo-cal stratigraphic ranges and can be useful in stratigraphic correlation and paleoenvironmental reconstruction (Fig. 4). As with trace fossils (Seilacher 1970) and stromatolites (Bertrand-Sarfati and Walter 1981), pedotypes do not need temporal and genetic continuity to be useful in stratigraphy and paleoecology. There is a general progression of differ-ent pedotypes in the Cenozoic of Oregon (Fig. 4), aside from a few repetitions, such as the return of green (Xaxus) pedotypes after brown (Plas) pedotypes in the Haystack Valley Member of the upper John Day Formation, and the return of brown clayey (Skwiskwi) pedotypes from the Big Basin Member of the John Day Formation in the Mascall and Rattlesnake Formations. Each formation has a unique sequence of pedotypes that reflects temporal changes in soil-
forming conditions.Diversity, overturn, appearances, and disappearances of
pedotypes through time can be estimated (Fig. 5) in the same way as evolutionary metrics of mammal fossils (Alroy et al. 2000, Janis et al. 2002). Unlike biological species, which persist from evolution to extinction, ecosystems and their soils can be dispersed and reconstituted as conditions allow (Bernabo and Webb 1977). Thus, Skwiskwi paleosols in the upper Big Basin Member (30 Ma) and the Mascall (15 Ma) and Rattlesnake formations (7 Ma) all represent subhumid woodland communi-ties, even though their taxonomic compositions were probably quite different. With these few exceptions, stratigraphic ranges defined by the first and last occurrences of pedotypes are comparable in duration with stratigraphic ranges of associated mammals (Alroy et al. 2000, Janis et al. 2002). In addition, both pedotype and mammalian diversity and turnover were generally greater during the warm-wet late Eocene (34–45 Ma) and middle Miocene (19–12 Ma) than during the cold-dry late Oligocene and late Miocene-Pleistocene (Alroy et al. 2000, Janis et al. 2000; Fig. 5).
Each pedotype also represents a particular kind of vegetation, geomorphic setting, parent material, and time for formation (see Retallack et al. 2000, 2002a, Retallack 2004a, b). The relative
Figure 2. Field photographs of selected paleosols: A. Ultisols and Alfisols above the Nut Beds near Clarno. B. Alfisols and Gleyed Incep-tisols in the central Painted Hills. C. Aridisols and Andisols at Foree north of Kimberly. D. Caprock Aridisols south of Kimberly.
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Figure 3. Chemical and petrographic data on selected Oregon paleosols. A. Xaxus and Xaxuspa paleosols (revised section) from Foree. B. Luca paleosol from Whitecap Knoll near Clarno. C. Lakayx and Scat paleosols above Nut Beds near Clarno (data from Re-tallack et al. 2000).
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importance of these paleoenvironmental factors through time can be assessed from the relative abundance of pedotypes for two million year bins centered on each million-year increment (Fig. 6). Forest pedotypes were common in the Eocene (35–45 Ma), woodland pedotypes in early Oligocene (30–35 Ma) and the middle Miocene (19–12 Ma), and desert shrubland (prob-ably sagebrush) pedotypes at other times. Forest and woodland paleosols have large root traces and differ in the degree of chemical and mineral weathering of the clayey subsurface (Bt) horizons (Retallack et al. 2000), whereas shrubland paleosols
Figure 4. Stratigraphic range of named pedotypes of paleosols from Oregon and Washington.
Figure 5. Richness (A), appearances (B), disappearances (C), overturn (D), calcareous paleosols or pedocals (E) and taxo-nomic orders (F) through time of pedotypes of paleosols from Oregon and Washington. Bin size is 1 million years either side of each million year increment. Data source is 1856 paleosols in 50 described pedotypes (Retallack et al. 2000, 2002a, Bestland and Forbes 2003, Retallack 2004a,b).
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have shallow calcareous nodules (Bk) and common cicada bur-rows (Retallack 2004a). There has also been a steady increase in bunch grasslands with granular-ped paleosols since the early Oligocene (30 Ma), and sod grasslands with crumb-ped pa-leosols since the early Miocene (19 Ma: Fig. 6A). Streamside habitats with very weakly developed pedotypes were most com-mon during the warm-wet Eocene-early Oligocene (>30 Ma) and middle Miocene (19–12 Ma), whereas chemically oxidized paleosols of well-drained floodplains were common at other times (Fig. 6B). Although coarse-grained fluvial sediments and volcanic lahars are common in the Eocene-early Oligocene and middle Miocene, most of the paleosol record was formed on claystones weathered from redeposited rhyodacitic airfall tuffs (Fig. 6C; Bestland 2000, Retallack et al. 2000).
Some quantifiable aspects of paleosols reflect vegetation and the activities of soil-dwelling animals. Modern soils of eastern Washington have large (1–2 cm in diameter) burrows of cica-das in sagebrush soils, abundant small (0.5 mm in diameter) earthworm pellets in grasslands, and medium to coarse (1–5 cm) blocky peds in forests (O’Geen and Busacca 2002). The fine crumb structure of sod grassland soils, granular structure of bunch grassland soils, and blocky structure of woodland soils can also be recognized in paleosols (Retallack 2004a,b). The decline in minimum size of peds in paleosol surface horizons over the past 45 Ma in Oregon (lower envelope to Fig. 7A) produces a corresponding increase in internal surface area of paleosols (upper envelope of Fig. 7B). Internal surface area was calculated from the number of approximating cubes to the average ped size in the thickness of A horizons, added to a comparable calculation for Bt or Bw horizons when present. These measurements support an early Oligocene (30 Ma) advent of bunch grassland, early Miocene (19 Ma) appearance of sod grassland, and late Miocene (7 Ma) appearance of tall sod grassland in Oregon from pedotypes (Fig. 6A). Despite these innovations, both average ped size and internal surface area have changed little over the past 45 Ma.
Soil Taxonomy
There are 12 orders in the soil taxonomy of the US Soil Conservation Service (Soil Survey Staff 1999). Identification of soils and paleosols in this classification system requires laboratory analyses (Fig. 2). For paleosols, there are petro-graphic and chemical proxies that accommodate alterations to soil during burial (Retallack 1997a, 2001a). Comparable laboratory data also is needed for identification of paleosols in other soil classifications, such as map units of the United Nations Educational Scientific and Cultural Organization’s soil maps of the world (Food and Agriculture Organiza-tion 1974), and traditional soil names (Stace et al. 1968). A simplified field classification (Mack et al. 1993) is useful for labelling paleosols, but not for paleoenvironmental in-terpretation (Retallack 1993, Mack and James 1994), which is emphasized in the present study.
Details of soil taxonomic determinations are published elsewhere (Retallack et al. 2000, 2002a, Retallack 2004a,b),
but a guide to key criteria is presented here. Many paleosols dominated by volcanic shards (Fig. 3A) have non-crystalline colloids that have been partly recrystallized to clinoptilolite and celadonite (Retallack et al. 2000, 2002a, Bestland 2002), as in Andisols (suffix “-and” of Soil Survey Staff 1999) and Andosols (Food and Agriculture Organization 1974). Some paleosols have the crumb texture, thin root traces, and thick-ness of this fine soil fabric to qualify as Mollisols (suffix “-oll” of Soil Survey Staff 1999) and Kastanozems (of Food and Agriculture Organization 1974). Other paleosols have high soda/potash ratios, and shallow nodules and concretions of calcite and chalcedony, which are characteristic of Aridisols (suffix “-id” of Soil Survey Staff 1999) and Xerosols, Solon-chak, and Solonetz (of Food and Agriculture Organization 1974). Paleosols that are weakly developed with bedding
Figure 6. Variation in soil forming factors of vegetation (A), geo-morphic setting (B), parent material (C) and degree of develop-ment (D) through time of pedotypes of paleosols from Oregon and Washington. Bin size and data source are same as for Figure 5.
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planes of sedimentary parent material little disrupted by root traces are classified as Entisols (suffix “-ent” of Soil Survey Staff 1999) and Fluvisols (Food and Agriculture Organization 1974). These can be contrasted with moderately developed paleosols in which all traces of sedimentary organization have been obliterated by the activity of worms and roots, and which have a subsurface horizon that has been enriched (by at least 8%) in clay washed down into the profiles, as in Alfisols (Fig. 3B; suffix “-alf” of Soil Survey Staff 1999) and Luvisols (Food and Agriculture Organization 1974). All of the paleosols noted above are rich in cationic nutrients (Ca2+, Mg2+, Na+, K+); however, the paleosols were significantly less fertile if the molar sum of these bases divided by molar alumina is less than 0.5 (Sheldon et al. 2002). Such base-depleted paleosols with a subsurface clay-enriched horizon can be regarded as Ultisols (Fig. 3A; suffix “-ult” of Soil Survey Staff 1999) and Acrisols (Food and Agriculture Organization 1974). Strongly base-depleted, uniformly clayey, and often ferruginous and bauxitic paleosols can be identified as Oxisols (suffix “-ox” of Soil Survey Staff 1999), Ferralsols, and Nitosols (Food and Agriculture Organization 1974, Bestland et al. 1996).
Taxonomic interpretationThe primary purpose for identifying paleosols within modern
soil taxonomies is to find modern soils analogous to paleosols. Fossil soils, like other fossils, have lost much of their original form and function during burial and other alterations (Retallack 1997a). In the same way that behavior and ecology of fossil snails and mammals can be inferred from their nearest living relatives, much can be learned by comparison of paleosols with comparable soils. Soil classification is a simple way of navigat-
ing the vast array of modern soil descriptions in county and regional soil surveys. The Food and Agriculture Organization (1974) soil maps of the world are especially useful in this regard because they are global in coverage. It is primarily taxonomic considerations that enable comparison of the suite of paleo-sols from the middle Eocene Clarno Nut Beds (Figs. 2A, 3A) with soils now forming under the tropical forest (“selva of Lauraceae”) on Volcán San Martin, Mexico, and the suite of paleosols from the middle Oligocene Turtle Cove Member of the John Day Formation (Figs. 2C, 3C) with soils now form-ing under deciduous grassy woodland near Tehuacán, Mexico (Retallack et al. 2000).
Also informative is the relative abundance through time of soil orders (Fig. 5F) and Marbut’s (1935) simpler classification of calcareous soils (pedocals) versus non-calcareous soils (pe-dalfers: Fig. 5E). The onset of pedocals in the early Oligocene (30 Ma) is striking, as is their rarity in the middle Miocene (19–12 Ma). Pedocals reflect dry and seasonal climates with insufficient moisture to leach carbonate from soils. Forest soils (Oxisols, Ultisols, Alfisols) are common in the Eocene and early Oligocene (30–45 Ma), and Alfisols reappear during the middle Miocene (19–12 Ma) when woodland and highly sea-sonal soils (Andisols and Vertisols) are common. Desert soils (Aridisols) have been common at other times, with a fluctuating but increasing abundance of grassland soils (Mollisols) since the early Miocene (19 Ma).
TRANSFER FUNCTIONS
Paleoenvironmental interpretations can be quantified for paleosols by applying the experimental designs of Jenny (1941). His space-for-climate approach to quantifying soil formation uses a collection of modern soils from different climates, but that are comparable in other soil-forming factors, including vegetation, parent material, topographic setting, and time for formation. From this set of soils (a climosequence) can be extracted mathematical relationships (climofunctions) between soil variables (such as base deple-tion) and climatic variables (such as mean annual precipita-tion). Similarly, biofunctions for vegetation or other biotic effects can be extracted from biosequences, topofunctions for geomorphic changes from toposequences, lithofunc-tions for parent material effects from lithosequences, and chronofunctions for time of soil development from chrono-sequences (Retallack 2005). Generally applicable biofunc-tions, topofunctions, lithofunctions and chronofunctions for paleosols have yet to be devised, and for those aspects of paleoenvironmental interpretation, pedotype and taxonomic approaches are most useful (Fig. 6).
In contrast, there are several climofunctions from which to choose (Sheldon et al. 2002, Retallack 2005). Several features of modern soils show significant relationships with mean an-nual precipitation (P in mm) and mean annual temperature (T in oC): (1) depth to calcareous nodules (D in cm) in the soil profiles (Retallack 2005); (2) chemical index of alteration without potash (C, which is the molar ratio of alumina over
Figure 7. Variability in (A) ped size (cm) within surface horizons and (B) in calculated internal surface area versus land surface area (mm2/mm2) in paleosols of Oregon and Washington through time. Each plotted point represents one paleosol.
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alumina plus lime, magnesia, and soda times 100) of paleosol subsurface (Bt or Bw) horizons (Sheldon et al. 2002), and (3) alkalies index (S which is molar potash plus soda over alumina) of paleosol subsurface (Bt or Bw) horizons (Sheldon et al. 2002) , according to the equations:
P = 139.6 + 6.39D + 0.013D2 , where S.E. = ± 141 mm; R2 = 0.62
P = 221.12e0.0197C, where S.E. = ± 182 mm; R2 = 0.72T = -18.52S + 17.3, where S.E. = ± 4.4oC; R2 = 0.37
Depth to calcareous nodules can be corrected for burial compaction (using Inceptisol physical constants of Sheldon and Retallack 2001), but atmospheric CO2 levels and erosion of the paleosols before burial were not considered sufficiently significant to require correction (Retallack et al. 2000, Retal-lack 2004a,b).
Climofunction application
When applied to calcic and chemical data from Oregon paleosols, each paleoclimatic transfer function indicates long term oscillations between arid-cool and humid-warm paleoclimate. The covariance of paleotemperature and paleo-precipitation is striking from chemical climofunctions (Fig. 8A). Paleoprecipitation estimated from depth to carbonate (Fig. 8B) is comparable with that estimated from chemical composition. There are limitations to carbonate data, because carbonate-bearing paleosols appeared in Oregon only after 30 Ma, and such paleosols cannot record high precipitation, which creates non-calcareous soils (Retallack 2005).
Some parts of Oregon’s fossil soil record show exceptional temporal resolution. During the late Oligocene (28.6–23.4 Ma) accumulation of the upper Turtle Cove Member on Longview Ranch near Kimberly there are 105 climatic fluctua-tions recorded by depth to carbonate in 358 paleosols, as well as by the oxygen and carbon isotopic composition of their pedogenic carbonate (Fig. 9). These variations principally record paleoclimatic cycles paced with changes in obliquity (41 k.yr) of the Earth’s orbit, a temporal resolution generally associated with Quaternary paleoclimatic records (Retallack et al. 2004).
Paleosol-based paleoclimatic curves are comparable with long-term variation in oxygen and carbon isotopic compo-sition of benthic marine foraminifera (Fig. 8D), and with variation in atmospheric CO2 estimated from stomatal index of Ginkgo leaves (Fig. 8E). Also comparable are time series of North American mammalian diversity, origination and extinction (Alroy et al. 2000, Janis et al. 2002), and diversity and floristic affinities of broadleaf forest vegetation in the lacustrine deposits of Oregon and Washington (Graham 1999). Paleosol data confirm Meyer and Manchester’s (1994) estimate of a mean annual precipitation of 1000-1500 mm and mean annual temperature of 3–9oC for the early Oligocene (32 Ma) Bridge Creek floras of central Oregon, and Chaney and Axelrod’s (1959) estimate of a mean annual precipitation of about 1270 mm and mean annual temperature of about 17oC for the middle Miocene (15 Ma) Mascall flora. Unlike
floral estimates, which require large collections and are scat-tered in time, paleosol data are easier to obtain and much more abundant through time.
An unexpected result, also found in comparable studies in Montana and Nebraska (Retallack 2007), is that paleoclimatic drying encouraged the spread of desert shrubland rather than grasslands (Fig. 6A). Grasslands appear for the first time and then spread at times of warm-wet climate (Figs. 6, 8) and high soilscape diversity (Fig. 5) in the early (30 Ma) Oligo-cene and early (19 Ma) and late (7 Ma) Miocene. Ancient grasslands are here inferred from paleosol data (Figs. 6, 7, 10), although there is a record of grass pollen (Leopold and
Figure 8. Cenozoic variation in mean annual precipitation (A) and mean annual temperature (B) inferred from paleosol chem-istry (A, B) and depth to calcic horizons (B), compared with oxygen and carbon isotopic composition of marine foraminifera (C–D), and CO2 levels inferred from stomatal index (E). Data of A and B from Retallack (2004a,b) and Retallack et al. (2000), with flanking curves one standard error from the transfer func-tion; data of C and D from Zachos et al. (2001); data of E from Retallack (2001b, 2002) with flanking curves one standard devia-tion of the stomatal index measurement, using transfer function of Wynn (2003).
98 PALEOBIOS, VOL. 28, NUMBER 3, JANUARY 2009
Denton 1987), phytoliths (Strömberg 2002, 2004, 2005, Blinnikov et al. 2002) and a few megafossils (Graham 1999, Retallack 2004b). Most of the fossil record of plants comes from swampy lowland deposits, whereas grasslands were mainly a phenomenon of dry, well drained soils (Retallack 1998). The evolution of grasslands can be inferred from root traces and aggregate size in paleosols. Soil aggregates (peds) are small in grassland soils because defined by a net-work of fine roots, and because produced as fecal pellets by earthworms, both characteristic of grassland soils (Retallack 1997a, 2001c). These highly coevolved ecosystems did not simply fill in arid inland regions, which were occupied then, as now, by desert shrublands.
Surprisingly, the uplifting of the Cascade, Klamath, and Olympic mountain ranges does not appear to correspond to the spread of rain-shadows, as traditionally inferred from fossil floras (Chaney 1948, Ashwill 1983), and more recently from declining isotopic values of oxygen in tooth enamel and pedogenic carbon (Kohn et al. 2002, Kohn and Fremd 2007). Growth of the Oregon Cascades inferred from radio-metrically dated eruptions (McBirney et al. 1974, Priest 1990) shows peaks of volcanic activity at 35, 25, 16, 10 and 4 Ma, but these are all times of warm-humid rather than cold-dry paleoclimate (Fig. 8A–B). Re-evaluation of the physiog-nomy of fossil floras from the northern Great Basin indicate high altitudes since at least the early Miocene, for example 2100 m for the middle Miocene (16 Ma) 49 Camp flora of northwestern Nevada (Wolfe et al. 1997), again at a time of warm-humid paleoclimate (Fig. 8A–B) and high atmospheric CO2 (Retallack 2001b, 2002). Eastern Oregon was broadly elevated during middle Miocene (16 Ma) initial eruptions of the Yellowstone hot spot, then subsided subsequently (Hum-phreys et al. 2000), with the exception of remarkably focused uplift of the Wallowa Mountains (Hales et al. 2006). Cascades and Klamath Mountain rain shadows have been important in creating generally dry long-term climate in eastern Oregon, with loessial sediments and calcareous paleosols starting during the Oligocene (30 Ma: Retallack et al. 2000). Rocky Mountain barriers also promoted summer-dry climate due to a cool nearby ocean by isolation of eastern Oregon from Gulf Coast cyclonic circulation starting during the early Miocene (19 Ma: Retallack 2004a). Orographic rain shadows were thus long-term boundary conditions to circulation since at least 50 Ma (Kent-Corson et al. 2006), but do not explain observed short-term paleoclimatic variation such as middle Miocene warm-wet spikes (Retallack 2007).
Also surprising are mismatches between broadly compa-rable paleoclimatic records of Oregon paleosols and oxygen and carbon isotopic composition of deep marine benthic foraminifera (Fig. 8C–D). Abrupt terminal Eocene, late Oligocene and Pleistocene change in oxygen isotopic data contrast with ramps in paleosol data (Sheldon et al. 2002, Sheldon and Retallack 2004), and the general trend of oxygen isotope records is sloping whereas paleosol records are level with deviations (Fig. 8C). Abrupt terminal Eocene, late Oli-
gocene and Pleistocene shifts in oxygen isotope values in the deep ocean are due to growth and decay of the Antarctic and Arctic ice caps, which abstracted large volumes of isotopically light oxygen from the ocean-atmosphere system (Zachos et al. 2001). Long term drift of marine oxygen isotopic values to lower values also is well known, and is attributed to crustal recycling (Veizer et al. 2000, Kasting et al. 2006) or burial diagenesis (Mii et al. 1997). Other theoretically plausible, but quantitatively unlikely reasons for mismatches of oceanic oxygen and Oregon paleosol records include evolution of C4 photosynthetic pathways in grasses which select CO2 heavy for either carbon or oxygen, and Rayleigh distillation to lower oxygen isotope values on land increasingly elevated and isolated from the sea by mountains (Retallack 2007). Less surprising is the carbon isotope value of the same deep marine foraminifera, which matches Oregon paleoclimatic records well (Fig. 8D), apart from Plio-Pleistocene deviation attributable to the rise of C4 vegetation on land (Cerling et al. 1997).
IMPLICATIONS FOR CENOZOIC GLOBAL CHANGE
The Cenozoic Oregon paleosol record is unusually long and detailed. Oregon’s back-arc paleotopography, rate of basinal subsidence, and rhyolitic parent materials changed little over the past 45 Ma, but Oregon’s changing climate and vegetation reflects a global paleoclimate signal for several reasons. Annual variation in oxygen isotopic composition of growth bands within Miocene mammalian tooth enamel reflect planetary revolutions (Fox 2001, Kohn et al. 2002). Also, temporal variations in depth to carbonate and its stable isotopic composition (Fig. 9) have the same periodicity (41-100 ka) as orbital variations in obliquity and eccentricity (Retallack et al. 2004). Late Eocene and middle Miocene peaks in temperature and precipitation in Oregon paleosols (Fig. 8A–C) correspond to highs in global atmospheric CO2 revealed by stomatal index of fossil Ginkgo (Fig. 8E) and peaks in deep ocean temperature inferred from marine oxygen isotopic values (Fig. 8D). Cenozoic global cooling from late Eocene (35 Ma) and middle Miocene (16 Ma) peaks of warmth has been explained by mountain uplift (Raymo and Ruddiman 1992), changing oceanic currents (Kennett 1982, Broecker 1989) and grassland coevolution (Retallack 2001c), and the Oregon paleosol record is relevant to each of these hypotheses.
Mountain uplift
Raymo and Ruddiman (1992) proposed that Himalayan and other mountain uplift promoted deep weathering by carbonic acid and was thus a sink for atmospheric carbon dioxide promoting Cenozoic climatic cooling after about 40 Ma. Their argument was based largely on the temporal in-crease of 87Sr/86Sr ratios, now attributed largely to carbonate weathering, which does not appreciably affect atmospheric redox, unlike silicate weathering (Jacobson et al. 2002, Ja-cobson and Blum 2003). There are additional fundamental
RETALLACK—CENOZOIC COOLING AND GRASSLAND EXPANSION IN OR AND WA 99
problems with Raymo and Ruddiman’s (1992) hypothesis be-cause metamorphic and volcanic emissions of carbon dioxide together with reduced chemical weathering at high altitude make mountain uplift a force not for global cooling, but for global warming (Zeitler et al. 2001). Himalayan and other mountain uplift promotes physical weathering, especially by means of glaciation, but not carbon-consuming chemical weathering of silicates (Jacobson and Blum 2003). Warm-wet spikes at 35 and 16 Ma in climate records of Oregon (Fig. 8) and elsewhere around the world (Retallack 2007) coincide with peaks in volcanic activity in the western US (McBirney et al. 1974, Priest 1990), and in globally significant eruptions of Ethiopia-Yemen and Columbia River flood basalts (Courtillot 2002). Some of these eruptions may have intruded coal (such as Oregon’s Herren Formation) and released to the atmo-sphere large amounts of thermogenic methane, which then oxidized to atmospheric CO2, as has been argued for other global warming events (Retallack and Jahren 2008). These warm-wet spikes also coincide with the Chesapeake impact crater at 35 Ma (Poag et al. 2003), and Ries and Steinheim craters at 16 Ma (Stoeffler et al. 2002), which would induce initial cooling from particulate ejecta in the atmosphere fol-
lowed by warming effects of carbon dioxide from destroyed soils and biomass (Poag et al. 2003).
In the Oregon paleosol record, which includes the Colum-bia River basalts (Sheldon 2003, 2006), the most chemically weathered paleosols are those least disturbed by volcanism, uplift and sedimentation (Bestland et al. 1996, Retallack et al. 2000). These paleosols would have consumed significant amounts of carbon dioxide by hydrolysis, especially at times of high atmospheric carbon dioxide and warm-wet weath-ering (Fig. 8). Thus, times of volcano-tectonic quiescence promote cooling, by consuming atmospheric carbon dioxide during hydrolytic weathering and fertilizing the ocean with bicarbonate and nutrients for deep oceanic storage of reduced carbon largely from phytoplankton (Retallack 2001c). In contrast, perturbations of mountain uplift, volcanism, and bolide impact are forces for planetary warming.
Ocean gateways
Ocean current reorganization as a result of continental drift has also been proposed as a force for Cenozoic global cooling. Kennett (1982) argued that cooling across the Eo-cene-Oligocene transition was caused by continental drift of
Figure 9. Milankovitch-scale fluctuations in mean annual paleoprecipitation inferred from depth to calcic (Bk) horizon within paleo-sols of the upper John Day Formation on Longview Ranch, near Kimberly.
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Australia and South America away from Antarctica, so that the ice continent was then thermally isolated by the Circum-Antarctic current. Broecker (1989) argued that Panamanian isthmus closure created a warm salty Caribbean and stimu-lated thermohaline circulation into North Atlantic bottom water formation to bring on the Pleistocene ice age. Problems with timing and mechanism have emerged for both ideas. The Circum-Antarctic current was initiated at 37 Ma rather than the 33.9 Ma Eocene-Oligocene boundary (Exon et al. 2004) and completed by 25 Ma (Lyle et al. 2007) when Oregon and other isotopic data indicate warming and deglaciation in a reversal of the proposed effect (Zachos et al. 2001). Pana-manian isthmus closure was at 2.8–3.1 Ma rather than the 1.8 Ma Plio-Pleistocene boundary (Bartoli et al. 2005), also a time of warming. More problematic is modeling indicating warming rather than cooling by diminution of the Humboldt Current as Chile drifted north (DeConto and Pollard 2003), and by Gulf Stream redistribution into the North Atlantic of warmth from an enclosed Caribbean Sea (Lund et al. 2006). Thermal transport modeling now suggests a more important role for global carbon dioxide than southern ocean currents in Cenozoic cooling (Huber and Nof 2007).
Southern and central American oceanic gateways are re-mote from Oregon records (Fig. 8A–B) which show general similarity with oceanic isotopic records (Fig. 8C–D) domi-nated by cores from the Southern Ocean (Zachos et al. 2001). Eocene-Oligocene cooling in the northern Pacific Ocean was synchronous with that in the southern hemisphere, yet not accompanied by changes in northern oceanic basin configura-tion (Scholl et al. 2003), lending support to climate-change mechanisms of global reach such as the greenhouse effect of atmospheric carbon dioxide.
Grassland coevolution
A third suggested cause of Cenozoic climatic cooling is stepwise coevolution of grassland grasses and grazers into novel ecosystems (sod grasslands) and soils (Mollisols) over the past 35 Ma (Retallack 2001a). Evidence for this view is well documented in Oregon (Fig. 10), where it has a long pedigree back to Thomas Huxley’s insights on horse evolu-tion from his 1876 examination of the collection of Oregon fossils O.C. Marsh at Yale obtained from Thomas Condon (Clark 1989). The coeval appearance of hypsodont horses and Mollisols in Oregon reflects grass-grazer coevolution that is also evident in Montana, Nebraska, Pakistan, and Kenya (Retallack 1995, 1997b, 2007, Retallack et al. 2002b). The appearance of open-country grass phytoliths 4 Ma.earlier than hypsodont horses in the Great Plains (Strömberg 2006) does not contradict this conclusion because the open country grass phytoliths come from desert and rangeland soils (Aridisols and Inceptisols) comparable with those now supporting sagebrush and bunch grassland (Retallack 2007). Other phy-toliths in these paleosols are considered by Strömberg (2004, 2006) as forest indicators, but they are indistinguishable from phytoliths of modern desert shrubs (Artemisia tridentata)
and small trees (Juniperus occidentalis) (see Blinnikov et al. 2002) and may also be from arid shrubland communities indicated by paleosols.
Considering the combined evidence from paleosols and phytoliths, grasslands did not merely adapt to climate change, but were a biological force for global change. Compared with preexisting woodland soils, grasses created soils richer in re-duced carbon and finer in structure deeper within the profile (Figs. 7, 10), thus promoting carbon sequestration in soils and derived sediments, as well as carbon dioxide consumption by hydrolytic weathering. Woodlands create moist air and dry soils by active transpiration, but grasslands create moist soil and dry air, which is cooler without greenhouse-inducing water vapor. Finally albedo of grasses is higher than that of woodlands, reflecting much solar warmth back into space (Retallack 2001c). Thus, grasslands and their soils (Molli-sols) became a force for planetary glaciation comparable with Late Devonian to Carboniferous (300-390 Ma) evolution of trees (Berner 1997) and soils of forests (Alfisols) and swamps (Histosols: Retallack 2001a, 2004c).
Unlike ocean gateways or mountain uplift, this mecha-nism of global cooling is biological not physicochemical. In the long term, weedy reproduction and abrasive phytoliths of grasses withstood the onslaught of increasingly large, hypsodont and hard-hooved mammals more effectively than pre-existing woody plants, but key stages in the process are evident in Oregon (Fig. 10). An early stage appearing by 35 Ma in Nebraska and 30 Ma in Oregon is marked by abundant opal phytoliths of grasses, cursorial and lophodont mixed browser-grazer rhinos and horses, and granular-structured calcareous paleosols (Retallack 1983, Retallack et al. 2000, Janis et al 2002, Strömberg 2002, 2004) A likely biological adaptation triggering the appearance of these fossils and paleosols is the origin of rumination (cud chewing) dated to 34–35 Ma by molecular clock techniques on digestive ribonucleases of ruminants (Jermann et al. 2001). Another coevolutionary advance at a time of warm-wet climate was appearance of hypsodont horses, abundant open country phytoliths and crumb-structured paleosols (Mollisols) of short sod grasslands in Oregon, Montana and Nebraska at about 19 Ma (early Miocene: Retallack 2007). The seminal evolutionary innovation for these communities may have been pack-hunting, indicated by the prorean gyrus in brain casts of fossil carnivores (Van Valkenburgh et al. 2003), which in turn promoted herding, intensive “cell grazing” and sod development (Savory and Butterfield 1999). A final coevolutionary advance, again at a time of warm-wet climate at about 7 Ma, was the appearance of large horses, along with deep-calcic, crumb structured paleosols (Mollisols) of subhumid climates and tall-grasslands (Retallack et al. 2002a). The seminal evolutionary adaptations in this case may have been monodactyl hard hooves and highly hypso-dont teeth, harder on seedlings of trees than grasses, so that the grassland-woodland ecotone rolled outward into more humid regions (Retallack 1997b). The 7 Ma coevolutionary
RETALLACK—CENOZOIC COOLING AND GRASSLAND EXPANSION IN OR AND WA 101
Figure 10. Coevolution of grasses and grazers. Although only examples of paleosols and fossil horses from Oregon are shown, compa-rable evolutionary change is known from all continents except Australia and Antarctica, and may have had global change consequences.
advances marked the onset of the C4 photosynthetic pathway as a global carbon isotopic perturbation (Cerling et al. 1997), although the pathway itself may be older; an origin of 40 Ma is estimated from paleosol carbon isotopic studies (Fox and Koch 2003), 25 Ma from phylogenetic analyses (Kellogg 1999), and 12.5 Ma from isotopic and anatomical studies of fossil grasses (Nambudiri et al. 1978, see Whistler and Bur-bank 1992, for revised age). There is as yet no evidence of C4 grasses in Oregon before their agricultural introduction, and such summer-dry climates have mainly C3 grasses today (Sage et al.1999). Cooling coevolutionary advances in carbon
consumption and sequestration, water vapor reduction and albedo increases of grasslands arrested tectono-volcanic and extraterrestrial warmings at 35, 16 and 7 Ma to create the bumpy ride of Oregon and global climate over the past 45 million years (Fig. 8).
CONCLUSIONS
The multifaceted Cenozoic paleosol record of Oregon and Washington undoubtedly has more to reveal about global change beyond the interpretations of changing vegetation and climate reviewed here. Especially promising is the role of
102 PALEOBIOS, VOL. 28, NUMBER 3, JANUARY 2009
past weathering in atmospheric carbon dioxide consumption (Bestland 2000) and mammalian community dynamics of different pedotypes (Retallack et al. 2004). New approaches to research can be expected beyond those of pedotypes, taxonomy, and transfer functions outlined here. Oregon and Washington’s Cenozoic paleosol record is continuous back to at least 45 Ma, and some parts of it are comparable in temporal resolution with deep-sea cores (Retallack et al. 2004). Comparable paleosol records are developing for Montana, Nebraska (Retallack 2007), and elsewhere (Retal-lack 2001a). Paleoenvironmental information from paleosols now supports evidence from fossil phytoliths and bones in paleosols at many localities, and provides new evidence for theories of global change.
ACKNOWLEDGMENTS
I thank Ted Fremd, Scott Foss, and Mary Oman for funding, research clearance, and curatorial assistance. Erick Bestland, Evelyn Krull, Jonathan Wynn, Megan Smith, Russell Hunt, and Nathan Sheldon helped with fieldwork and fossil collection. Ken Finger and Randall Irmis greatly improved the manuscript by diligent editing. Work was funded by National Park Service Contracts CX-9000-1-1009, 144-PX9345-99-005, and P1325010503, Bureau of Land Management cooperative agreement 8/21/01, and National Science Foundation grant EAR-0000953.
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App
endi
x 1.
Dia
gnos
es a
nd id
entifi
catio
ns o
f C
enoz
oic
pedo
type
s of
Ore
gon
and
Was
hing
ton.
Sah
aptin
mea
ning
is a
fter
Rig
sby
(196
5) a
nd D
elan
cey
et a
l. (1
988)
; so
me
pale
osol
s ar
e de
scri
bed
in t
his
pape
r, a
nd t
he o
ther
s ar
e de
scri
bed
by R
etal
lack
et
al. (2
000,
200
2a).
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otyp
e Sa
hapt
in
Ort
hogr
aphy
T
ype
profi
le
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gnos
is
U.S
. F.
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. m
ap
Aus
tral
ian
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axi
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er r
oot
Abi
áXi
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e C
reek
H
igh
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cla
yey
sand
ston
e w
ith
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eous
H
aplo
salid
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rthi
c So
lonc
hak
rhiz
ocon
cret
ions
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d re
lict
bedd
ing
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lonc
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e Á
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cock
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ple,
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ided
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face
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thic
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rric
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isol
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ater
itic
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larn
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hori
zon
wit
h re
d no
dule
s H
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hum
ult
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dzol
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in
Apá
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k su
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ndic
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thy
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Hill
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oriz
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rum
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ce (
mol
lic)
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p C
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nbod
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cm
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lcar
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rou
nd
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rum
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xtur
ed s
urfa
ce (
mol
lic)
over
sha
llow
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alci
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ll O
rthi
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nbod
en
(<50
cm
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lcar
eous
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yzem
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asca
ll R
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t be
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d Fl
uven
t E
utri
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l
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Bla
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on
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ne w
ith
loca
l H
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t E
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c H
isto
sol
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ic G
ley
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woo
d Il
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s Pic
ture
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ge
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ck, re
d (1
0R),
wit
h cl
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sub
surf
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Hap
luda
lf O
rthi
c L
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ol
Bro
wn
Podz
olic
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rizo
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bas
altic
rub
ble
and
basa
lt
Isci
t pa
th
Íš_í
t B
one
Cre
ek
Cru
mb-
stru
ctur
ed s
urfa
ce (
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over
sili
ceou
s D
urix
eral
f M
ollic
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row
n H
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duri
pan
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So
lonc
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So
il
Kal
as
raco
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Ka.
lás
Mas
call
Ran
ch
Cru
mb-
stru
ctur
ed c
laye
y (m
ollic
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rfac
e Pla
caqu
and
Mol
lic G
leys
ol
Wie
senb
oden
ov
er b
lack
-spo
tted
and
vei
ned
(pla
cic
horizo
n)
Ksk
us
smal
l K
skús
C
entr
al P
aint
ed
Thi
n re
ddis
h br
own,
roo
ted
surf
ace
Fluv
ent
Dys
tric
Flu
viso
l A
lluvi
al S
oil
H
ills
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alk
long
kw
a.lk
Pic
ture
Gor
ge
Car
bona
ceou
s cl
ayey
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) on
gra
y sh
ale
Aqu
ent
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yic
Luv
isol
H
umic
Gle
y
an
d ba
salt
Lak
ayx
shin
e L
a.kà
yX
Red
Hill
(C
larn
o)
Red
(2.
5YR
-10R
), c
laye
y, t
hick
, ab
unda
nt
Hap
ludu
lt Fe
rric
Acr
isol
K
rasn
ozem
sl
icke
nsid
ed c
utan
s, d
eepl
y w
eath
ered
Lak
im
soot
L
a.k’
im
Cen
tral
Pai
nted
B
row
n to
oliv
e, w
ith
root
tra
ces
and
blac
k A
quan
dic
Eut
ric
Gle
ysol
H
umic
Gle
y
Hill
s iro
n-m
anga
nese
nod
ules
Pla
caqu
ept
Luc
a re
d L
u_á
Whi
teca
p K
noll
Red
(2.
5YR
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), t
hick
, cl
ayey
, w
ith
Hap
luda
lf C
hrom
ic
Red
Pod
zolic
(Cla
rno)
re
mna
nt w
eath
erab
le m
iner
als
L
uvis
ol
Luq
uem
de
caye
d fir
e L
uq’_
m
Cla
rno
fern
qua
rry
Whi
te b
edde
d vo
lcan
ic a
sh w
ith
root
tra
ces
Typ
ic U
divi
tran
d V
itric
And
osol
A
lluvi
al
ÁÁ
RETALLACK—CENOZOIC COOLING AND GRASSLAND EXPANSION IN OR AND WA 107
App
endi
x 1.
(co
nt.)
Dia
gnos
es a
nd id
entifi
catio
ns o
f C
enoz
oic
pedo
type
s of
Ore
gon
and
Was
hing
ton.
Ped
otyp
e Sa
hapt
in
Ort
hogr
aphy
T
ype
profi
le
Dia
gnos
is
U.S
. F.
A.O
. m
ap
Aus
tral
ian
Maq
as
yello
w, or
ange
M
aqáš
C
arro
ll R
im
Bro
wn,
fine
blo
cky
peds
, th
ick
subs
urfa
ce
Vitr
ic
Och
ric
And
osol
N
on-c
alci
c So
il
ho
rizo
n H
aplu
stan
d
Mic
ay
root
M
í_ay
C
larn
o m
amm
al
Bro
wn
to o
live
clay
wit
h ro
ot t
race
s an
d A
quan
dic
Eut
ric
Fluv
isol
A
lluvi
al S
oil
qu
arry
re
lict
bedd
ing
Fluv
aque
nt
Mon
ana
unde
r-ne
ath
Mon
a.na
B
one
Cre
ek
Thi
n im
pure
lign
ite (
O)
on s
ands
tone
and
Sa
pris
t E
utri
c H
isto
sol
Aci
d Pe
at
silts
tone
wit
h sh
allo
w r
oot
trac
es
Nix
go
od
Níx
K
ahlo
tus
Thi
ck (
>50
cm)
crum
b te
xtur
ed (
mol
lic A
) H
aplo
xero
ll K
asta
noze
m
Pra
irie
Soi
l
ov
er d
eep
calc
areo
us lo
ess
Nuk
ut
flesh
N
ukút
B
row
n G
rott
o R
ed s
licke
nsid
ed c
lay
over
thi
ck v
iole
t L
ithi
c O
rthi
c A
cris
ol
Kra
snoz
em
(P
aint
ed H
ills)
sa
prol
ite o
n rh
yolit
e flo
w
Kan
hapl
udul
t
Nuq
was
th
roat
N
uq’w
as
Pic
ture
Gor
ge
Ora
nge
mod
erat
ely
wea
ther
ed b
asal
tic
Dys
troc
hrep
t D
ystr
ic
Non
-cal
cic
rubb
le o
ver
basa
lt
Cam
biso
l B
row
n So
il
Pasc
t cl
oud
Páš_
t R
ed H
ill (
Cla
rno)
O
live
gray
to
oran
ge, sl
icke
nsid
ed,
Som
brih
umul
t H
umic
Acr
isol
G
rey
Bro
wn
subs
urfa
ce c
laye
y (B
t) h
oriz
on
Podz
olic
Pata
t tr
ee
Páta
t H
anco
ck C
anyo
n B
edde
d sa
ndst
one
wit
h ro
ot t
race
s, m
ildly
Psa
mm
entic
E
utri
c B
row
n E
arth
(Cla
rno)
le
ache
d an
d fe
rrug
iniz
ed
Eut
roch
rept
C
ambi
sol
Patu
m
ount
ain
Pátu
B
one
Cre
ek
Cru
mb-
stru
ctur
ed b
row
n (2
.5Y)
sur
face
(A
) C
alci
xero
ll C
alci
c C
hern
ozem
ov
er s
hallo
w (
<50
cm)
mic
rite
-cha
lced
ony
K
asta
noze
m
rhiz
ocon
cret
ions
(B
k)
Pla
s w
hite
Plá
š B
one
Cre
ek
Whi
te, w
eakl
y st
ruct
ured
silt
ston
e w
ith
Ust
ic
Cal
cic
Xer
osol
G
rey
brow
n
de
ep (
>45
cm)
calc
areo
us n
odul
es
Hap
loca
lcid
Cal
care
ous
Soil
Pla
spa
In w
hite
Plá
š-pa
B
one
Cre
ek
Whi
te, w
eakl
y st
ruct
ured
silt
ston
e w
ith
Typ
ic
Cal
cic
Gre
y B
row
n
sh
allo
w (
<45
cm)
calc
areo
us n
odul
es
Hap
loca
lcid
Ye
rmos
ol
Cal
care
ous
Soil
Psw
a st
one,
cla
y Pšw
á be
low
Bla
ck
Pur
ple
clay
ey s
ubsu
rfac
e (B
t) w
ith
Lit
hic
Hap
luda
lf O
rthi
c L
uvis
ol
Bro
wn
Podz
olic
Spur
(C
larn
o)
core
ston
es a
nd g
rada
tiona
l con
tact
dow
n to
an
desi
te
Sak
onio
n Šá
.k
sout
hern
Pai
nted
R
ed, cl
ayey
, th
ick
sapr
olite
wit
h an
desi
te
Typ
ic P
aleu
dult
Ferr
ic A
cris
ol
Bro
wn
Podz
olic
Hill
s co
rest
ones
Saya
yk
sand
Sa
yáyk
w
Cla
rno
Nut
Bed
s B
edde
d si
ltsto
ne w
ith
root
tra
ces
Tro
poflu
vent
E
utri
c Fl
uvis
ol
Allu
vial
Scat
da
rk
S_át
be
low
Cla
rno
Thi
n gr
ay c
lays
tone
on
ande
sitic
H
aplu
mbr
ept
Hum
icl
Alp
ine
Hum
us
N
ut B
eds
cong
lom
erat
e
Cam
biso
l
108 PALEOBIOS, VOL. 28, NUMBER 3, JANUARY 2009
App
endi
x 1.
(co
nt.)
Dia
gnos
es a
nd id
entifi
catio
ns o
f C
enoz
oic
pedo
type
s of
Ore
gon
and
Was
hing
ton.
Ped
otyp
e Sa
hapt
in
Ort
hogr
aphy
T
ype
profi
le
Dia
gnos
is
U.S
. F.
A.O
. m
ap
Aus
tral
ian
Sita
xs
liver
Ši
t’áX
š R
ed H
ill (
Cla
rno)
O
live-
purp
le s
ilty
clay
ston
e, w
ith
relic
t Pla
caqu
and
Eut
ric
Gle
ysol
H
umic
Gle
y
be
ddin
g, a
nd ir
on-m
anga
nese
ski
ns a
nd
nodu
les
Skaw
sc
are
Skaw
Pic
ture
Gor
ge
Thi
n, d
ark
gray
to
blac
k sp
otte
d silts
tone
with
M
ollic
E
utri
c G
leys
ol
Hum
ic G
ley
root
tra
ces
over
bed
ded
tuff
aceo
us s
iltst
one
End
oaqu
ent
Skw
i-sk
wi
brow
n Sk
w’i-
skw
’i ce
ntra
l Pai
nted
R
eddi
sh b
row
n (7
.5YR
-10Y
R)
with
sub
surf
ace
Eut
ric
Och
ric
Bro
wn
Ear
th
H
ills
clay
ey (
Bt)
hor
izon
Fu
lvud
and
And
osol
Tat
as
bask
et
Tát
aš
Mas
call
Ran
ch
Cru
mb-
stru
ctur
ed, br
own
clay
ey s
ilty
surf
ace
Cal
ciud
oll
Cal
cic
Che
rnoz
em
over
dee
p (>
50 c
m)
calc
areo
us n
odul
es a
nd
Kas
tano
zem
ca
lcar
eous
rhi
zoco
ncre
tions
Tic
am
eart
h ti.
_ám
ce
ntra
l Pai
nted
R
ed (
7.5Y
R-5
YR),
wea
k su
bsur
face
A
ndic
E
utri
c C
hoco
late
Soi
l
Hill
s cl
ayey
(B
t) h
oriz
on
Eut
roch
rept
C
ambi
sol
Tili
wal
bl
ood
Tilí
wal
B
row
n G
rott
o V
ivid
red
(10
YR-7
.5YR
) cl
ayst
one
brec
cia
Plin
thic
Plin
thic
K
rasn
ozem
(Pai
nted
Hill
s)
wit
h dr
ab-h
aloe
d ro
ot t
race
s K
andi
udox
Fe
rral
sol
Tim
a w
rite
T
ima
Bon
e C
reek
G
ranu
lar
stru
ctur
e su
rfac
e (A
) ov
er c
laye
y N
atri
c D
urix
eral
f M
ollic
So
lone
tz
subs
urfa
ce (
Bt)
and
sili
ceou
s du
ripa
n (B
q)
So
lone
tz
Tla
l ci
cada
T
lal
Kah
lotu
s W
hite
loes
sic
with
abu
ndan
t sh
allo
w (
<50
cm)
Hap
loca
lcid
C
alci
c G
rey
Bro
wn
calc
areo
us n
odul
es a
nd a
bund
ant
cica
da
X
eros
ol
Cal
care
ous
Soil
burr
ows
Tna
n cl
iff
tnán
M
asca
ll R
anch
R
eddi
sh b
row
n (1
0YR
-7.5
YR),
cla
yey
Mol
lic
Cal
cic
Bro
wn
silty
, gr
anul
ar-c
rum
b st
ruct
ure
surf
ace
Hap
loca
lcid
X
eros
ol
Har
dpan
Soi
l
ov
er s
hallo
w (
<50
cm)
calc
areo
us
rhiz
ocon
cret
ions
Tuk
say
cup,
pot
T
uksá
y so
uthe
rn P
aint
ed
Red
(2.
5YR
-10R
), k
aolin
itic,
sub
surf
ace
Plin
thic
Plin
thic
R
ed P
odzo
lic
H
ills
clay
ey (
Bt)
hor
izon
, w
ith
rede
posi
ted
Pale
udul
t A
cris
ol
soil
clas
ts
Tut
anik
ha
ir
Tút
anik
U
nity
N
ear-
mol
lic b
row
n th
ick
surf
ace
(A)
Typ
ic
Hap
lic
Bro
wn
Ear
th
over
dee
p ca
lcar
eous
and
sili
ceou
s N
atri
xera
lf K
asta
noze
m
rhiz
ocon
cret
ions
Wal
asx
pine
gum
W
alá.
sX
Mas
call
Ran
ch
Pale
yel
low
cla
yey
(A)
wit
h sl
icke
nsid
ed
Chr
omud
ert
Chr
omic
B
row
n C
lay
clay
ski
ns
V
ertis
ol
RETALLACK—CENOZOIC COOLING AND GRASSLAND EXPANSION IN OR AND WA 109
110 PALEOBIOS, VOL. 28, NUMBER 3, JANUARY 2009
App
endi
x 1.
(co
nt.)
Dia
gnos
es a
nd id
entifi
catio
ns o
f C
enoz
oic
pedo
type
s of
Ore
gon
and
Was
hing
ton.
Ped
otyp
e Sa
hapt
in
Ort
hogr
aphy
T
ype
profi
le
Dia
gnos
is
U.S
. F.
A.O
. m
ap
Aus
tral
ian
Waw
cak
split
W
awc’
ak
sout
hern
Pai
nted
O
live
brow
n, c
laye
y, w
ith
defo
rmed
cla
y E
ntic
C
hrom
ic
Gre
y C
lay
H
ills
skin
s an
d pr
omin
ent
clas
tic d
ikes
C
hrom
uder
t V
ertis
ol
Yanw
a w
eak,
poo
r Ya
nwá
sout
hern
Pai
nted
T
hin,
bro
wn,
cla
yey
to s
ilty,
impu
re
His
tic
Hum
ic
Hum
ic G
ley
H
ills
ligni
tes
on a
roo
ted
unde
rcla
y H
umaq
uept
G
leys
ol
Yapa
s gr
ease
Yá
paš
Car
roll
Rim
D
ark
brow
n, fi
ne b
lock
y pe
ds (
A, B
w),
M
ollic
M
ollic
Pra
irie
Soi
l
(Pai
nted
Hill
s)
over
dee
p (>
50 c
m)
calc
areo
us
Hap
lust
and
And
osol
no
dule
s (B
k)
Yapa
spa
in g
reas
e Yá
paš-
pa
Bon
e C
reek
D
ark
brow
n, fi
ne b
lock
y pe
ds (
A, B
w),
V
itran
dic
Cal
cic
Gre
y B
row
n
ov
er s
hallo
w (
<50
cm)
calc
areo
us
Hap
loca
lcid
X
eros
ol
Cal
care
ous
Soil
nodu
les
(Bk)
Xau
s ro
ot
Xau
S M
asca
ll R
anch
B
row
n vo
lcan
icla
stic
san
dsto
ne w
ith
Cal
cari
c Psa
mm
ent
Allu
vial
Soi
l
re
lict
bedd
ing
and
root
tra
ces
Fluv
isol
Xax
us
gree
n X
áXuš
Fo
ree
Gre
en, fin
e bl
ocky
ped
s (A
, B
w),
ove
r A
quic
V
itric
W
iese
n-bo
den
(n
ear
Day
ville
) de
ep (
>50
cm)
calc
areo
us n
odul
es (
Bk)
U
stiv
itran
d A
ndos
ol
Xax
uspa
in
gre
en
XáX
uš-p
a Fo
ree
Gre
en, fin
e bl
ocky
ped
s (A
, B
w),
ove
r A
quic
C
alca
ric
Gre
y B
row
n
(nea
r D
ayvi
lle)
shal
low
(<5
0 cm
) ca
lcar
eous
nod
ules
(B
k)
Hap
loca
lcid
G
leys
ol
Cal
care
ous
Soil
App
endi
x 2.
Key
to
Cen
ozoi
c pa
leos
ols
of c
entr
al O
rego
n an
d W
ashi
ngto
n.
P
aren
t M
ater
ial
Subs
urfa
ce H
oriz
on
Surf
ace
Hor
izon
O
ther
Dif
fere
ntia
e Ped
otyp
e
1. R
hyod
aciti
c lo
essi
c 1.
1. R
elic
t be
ddin
g an
d ro
ot
1.1.
1. C
lear
rel
ict
bedd
ing
1.1.
1.1.
Gra
y, g
reen
or
brow
n cl
ayst
one
Mic
ay
silty
tuf
f tr
aces
onl
y
1.
1.1.
2. G
ray
to b
row
n si
ltsto
ne
Saya
yk
1.
1.1.
3. R
ed c
lays
tone
K
skus
1.
1.1.
4. B
row
n cl
ayey
silt
ston
e C
mti
1.
1.1.
5. W
hite
vol
cani
c as
h L
uque
m
1.
1.1.
6. O
live
brow
n cl
ay, w
ith
defo
rmed
cla
y sk
ins
Waw
cak
an
d pr
omin
ent
clas
tic d
ikes
1.
1.1.
7. Y
ello
w c
lay
wit
h de
form
ed c
lay
skin
s W
alas
x
1.1.
2. P
eat,
lign
ite o
r co
al
1.1.
2.1.
Bla
ck c
oal o
n cl
ayst
one
Cm
uk
1.
1.2.
2. B
row
n cl
ayey
lign
ite o
n si
ltsto
ne
Yanw
a
1.
1.2.
3. B
lack
cla
yey
ligni
te o
n sa
ndst
one
and
silts
tone
M
onan
a
1.
2. L
arge
bla
ck ir
on-m
anga
nese
1.
2.1.
Gra
y or
gre
en c
lays
tone
1.
2.1.
1. R
elic
t be
ddin
g L
akim
mot
tles
1.
3. S
mal
l dar
k gr
ay to
blac
k iro
n 1.
3.1.
Pla
ty t
o bl
ocky
ped
s
Skaw
man
gane
se n
odul
es
1.3.
2. C
rum
b te
xtur
ed s
urfa
ce
K
alas
(m
ollic
)
1.
4. C
lay
enri
ched
(ar
gilli
c),
1.4.
1. P
laty
to
bloc
ky p
eds
1.4.
1.1.
Pur
ple,
slic
kens
ided
, su
bsur
face
cla
yey
(Bt)
A
cas
no
n-ca
lcar
eous
, bl
ocky
ped
s (n
on m
ollic
) ho
rizo
n w
ith
red
nodu
les
1.
4.1.
2. R
ed (
2.5Y
R-1
0R),
sm
ectit
ic, w
ith
rem
nant
L
uca
w
eath
erab
le m
iner
als
1.
4.1.
3. R
ed (
2.5Y
R-1
0R),
kao
liniti
c, w
ith
Tuk
say
re
depo
site
d so
il cl
asts
1.
4.1.
4. V
ivid
red
(10
R-7
.5R
) cl
ayst
one
brec
cia
wit
h T
iliw
al
dr
ab-h
aloe
d ro
ot t
race
s
1.
4.1.
5. R
ed (
2.5Y
R-1
0R)
wit
h sl
icke
nsid
ed s
mec
titic
L
akay
x
clay
and
rar
e w
eath
erab
le m
iner
als
1.
5. C
laye
y en
rich
ed (
argi
llic)
1.
5.1.
Fin
e bl
ocky
to
gran
ular
1.
5.1.
1. N
on-n
odul
ar
Skw
iskw
i
and
gran
ular
str
uctu
red
stru
ctur
ed s
urfa
ce (
near
mol
lic)
1.
5.1.
2. d
eep
calc
areo
us a
nd s
ilice
ous
rhiz
ocon
cret
ions
T
utan
ik
RETALLACK—CENOZOIC COOLING AND GRASSLAND EXPANSION IN OR AND WA 111
App
endi
x 2.
(co
nt.)
Key
to
Cen
ozoi
c pa
leos
ols
of c
entr
al O
rego
n an
d W
ashi
ngto
n.
P
aren
t M
ater
ial
Subs
urfa
ce H
oriz
on
Surf
ace
Hor
izon
O
ther
Dif
fere
ntia
e Ped
otyp
e
1.
5.1.
3. S
ilice
ous
duri
pan
(Bq)
at
dept
h T
ima
1.5.
2. C
rum
b te
xtur
ed s
urfa
ce
N
ix
(mol
lic)
1.
6. C
lay
enri
ched
, bu
t no
t 1.
6.1.
Pla
ty t
o bl
ocky
str
uctu
red
1.6.
1.1.
Bro
wn
(2.5
Y-10
YR)
Maq
as
qu
ite a
rgill
ic, no
ncal
care
ous
surf
ace
(non
mol
lic)
1.
6.1.
2. R
ed (
7.5Y
R-5
YR)
Tic
am
1.
7 W
ell d
evel
oped
cal
care
ous
1.7.
1. M
assi
ve, pl
aty
or b
lock
y 1.
7.1.
1. W
hite
, tu
ffac
eous
Pla
s
nodu
les
deep
er t
han
50 c
m
peds
(no
n-m
ollic
)
1.
7.1.
2. G
reen
tuf
face
ous,
X
axus
1.
7.1.
3. B
row
n, t
uffa
ceou
s Ya
pas
1.7.
2. C
rum
b te
xtur
ed s
urfa
ce
C
il
(m
ollic
)
1.
8. W
ell d
evel
oped
cal
care
ous
1.8.
1 M
assi
ve, pl
aty
or b
lock
y 1.
8.1.
1. G
reen
tuf
face
ous
X
axus
pa
no
dule
s sh
allo
wer
tha
n 50
cm
pe
ds (
non-
mol
lic)
1.
8.1.
2. B
row
n tu
ffac
eous
Ya
pasp
a
1.
8.1.
3. W
hite
loes
sic
wit
h ab
unda
nt c
icad
a bu
rrow
s T
lal
1.1.
2. C
rum
b te
xtur
ed s
urfa
ce
C
ilba
(mol
lic)
1.
9. W
ell d
evel
oped
cal
care
ous
1.9.
1. M
assi
ve, pl
aty
or b
lock
y 1.
9.1.
1. W
hite
, tu
ffac
eous
,
Pla
spba
nodu
les
shal
low
er t
han
45 c
m
peds
(no
n-m
ollic
)
1.
10. Si
licic
rhi
zoco
ncre
tions
1.
10.1
. C
rum
b te
xtur
ed s
urfa
ce
1.10
.1.1
. C
alci
c ho
rizo
n de
eper
tha
n 50
cm
T
atas
and
less
er c
arbo
nate
at
dept
h (m
ollic
)
1.
10.1
.2. C
alci
c ho
rizo
n sh
allo
wer
tha
n 50
cm
Pa
tu
1.
11. T
hick
sili
ceou
s du
ripa
n
Is
cit
(B
q) a
t de
pth
2. V
olca
nicl
astic
san
dsto
ne
2.1.
Cla
yey
root
tra
ces
little
2.
1.1.
Rel
ict
bedd
ed s
ands
tone
2.
1.1.
1. L
ow-s
oda
sand
ston
e X
aus
or c
ongl
omer
ate
wea
ther
ed
2.
2. S
ilice
ous
rhiz
ocon
cret
ions
2.
2.1.
Rel
ict
bedd
ed s
ands
tone
2.
2.1.
1. H
igh
soda
cla
yey
sand
ston
e A
biax
i
little
wea
ther
ed
2.
3. M
ildly
leac
hed
and
2.3.
1. R
elic
t be
dded
san
dsto
ne
2.3.
1.1.
On
ande
sitic
con
glom
erat
e Pa
tat
fe
rrug
iniz
ed
112 PALEOBIOS, VOL. 28, NUMBER 3, JANUARY 2009
App
endi
x 2.
(co
nt.)
Key
to
Cen
ozoi
c pa
leos
ols
of c
entr
al O
rego
n an
d W
ashi
ngto
n.
P
aren
t M
ater
ial
Subs
urfa
ce H
oriz
on
Surf
ace
Hor
izon
O
ther
Dif
fere
ntia
e Ped
otyp
e
2.
4. L
ittle
wea
ther
ed
2.4.
1. T
hin
gray
cla
ysto
ne
2.4.
1.1.
On
ande
sitic
con
glom
erat
e Sc
at
2.
5. R
elic
t be
ddin
g, a
nd ir
on-
2.5.
1. O
live-
purp
le s
ilty
clay
ston
e 2.
5.1.
1. O
n an
desi
tic c
ongl
omer
ate
Sita
xs
m
anga
nese
cla
y-sk
ins
and
nodu
les
2.
6. W
ith
wea
k su
bsur
face
2.
6.1.
Bro
wn
pedo
lithi
c 2.
6.1.
1. O
n br
own
pedo
lithi
c co
nglo
mer
ate
Apa
x
clay
ey (
Bt)
hor
izon
co
nglo
mer
ate
2.
7. O
live
gray
to
oran
ge,
2.7.
1. O
live
gray
to
oran
ge
2.7.
1.1.
On
ande
sitic
con
glom
erat
e Pa
sct
slic
kens
ided
, cla
yey
(Bt)
hor
izon
cl
ayst
one
3. R
hyol
ite fl
ow
3.1.
Red
slic
kens
ided
cla
y 3.
1.1.
Red
slic
kens
ided
cla
y 3.
1.1.
1. T
hick
vio
let
sapr
olite
N
ukut
4. B
asal
tic fl
ow
4.1.
Thi
ck, re
d (1
0R),
cla
yey
4.1.
1. R
ed (
10R
) pl
aty
to b
lock
y
Iluk
as
su
bsur
face
hor
izon
4.
2. G
ray
shal
ey
4.2.
1. C
arbo
nace
ous
clay
ey
K
wal
k
4.
3. O
rang
e cl
ayey
4.
3.1.
Ora
nge
clay
ey
4.3.
1.1.
Ora
nge-
wea
ther
ed b
asal
tic r
ubbl
e N
uqw
as
5. A
ndes
itic
flow
5.
1. R
ed, c
laye
y, s
ubsu
rfac
e (B
t)
5.1.
1. R
ed c
lays
tone
5.
1.1.
1. C
ores
tone
s an
d th
ick
sapr
olite
Sa
k
6. A
ndes
itic
intr
usiv
e 5.
2. P
urpl
e cl
ayey
sub
surf
ace
(Bt)
5.
2.1.
Pur
ple
plat
y cl
ayst
one
5.2.
1.1.
Cor
esto
nes
and
thic
k sa
prol
ite
Psw
a
RETALLACK—CENOZOIC COOLING AND GRASSLAND EXPANSION IN OR AND WA 113