Topographic and climatic influences on accelerated loess accumulation since the last

glacial maximum in the Palouse, Pacific Northwest, USA

*Mark R. Sweeney, Department of Geology, Washington State University, Pullman, WA,

99164-2812

Alan J. Busacca, Department of Crop and Soil Sciences, Washington State University, Pullman,

WA 99164-6420

David R. Gaylord, Department of Geology, Washington State University, Pullman, WA, 99164-

2812

*Corresponding author. Current address: Desert Research Institute, 2215 Raggio Parkway, Reno,

NV 89512-1095. 775-673-7412, mark.sweeney@dri.edu

2

Topographic and climatic influences on accelerated loess accumulation since the last glacial

maximum in the Palouse, Pacific Northwest, USA

Abstract

Topographic and climatic influences have controlled thick loess accumulation at the

southern margin of the Palouse loess in northern Oregon. Juniper and Cold Springs Canyons,

located on the upwind flank of the Horse Heaven Hills, are oriented perpendicular to prevailing

southwesterly winds. These canyons are topographic traps that separate eolian sand on the

upwind side from thick accumulations (nearly 8 m) of latest Pleistocene to Holocene L1 loess on

the downwind side. Silt- and sand-rich glacial outburst flood sediment in the Umatilla Basin is

the source of eolian sand and loess for the region. Sediment from this basin also contributes to

loess accumulations across much of the Columbia Plateau to the northeast. Downwind of Cold

Springs Canyon, Mt. St. Helens set S and Glacier Peak tephras bracket 4 m of loess,

demonstrating that approximately 2500 g m-2

yr-1

of loess accumulated between about 15,400-

13,100 yrs B.P. Mass accumulation rates decreased to approximately 250 g m-2

yr-1

from 13,100

yr B.P. to the present. Tephrochronology suggests that the bulk of near-source Palouse loess

accumulated in one punctuated interval in the latest Pleistocene characterized by a dry and windy

climate.

Introduction

Quaternary loess and intercalated paleosols are the products of contrasting climatic

conditions (Kemp, 2001; Muhs and Bettis, 2003). Thick loess accumulations commonly are

linked to arid and windy glacial conditions marked by abundant sediment and sparse vegetation

in source areas that combine to produce high mass accumulation rates (MARs) (Roberts et al.,

3

2003; Antoine et al., 2001). Paleosols within many loess deposits developed during interglacial

or interstadial episodes and are typically associated with wetter conditions, deficient source

sediment, and low MARs (Kemp, 2001). Pedogenesis occurs to some degree throughout loess

accumulation, even when accumulation rates are high (Kemp et al., 1995; Kemp, 2001). Other

factors controlling loess accumulation include proximity to source, influence of eolian sand, and

interactions with topography (Mason et al., 1999; Pye, 1995). Separating climatic and non-

climatic factors that control loess accumulation is important when interpreting paleoclimatic

signals from loess.

Saltating sand grains play a key role in the entrainment of dust particles and the

formation of loess. Silt- and clay-sized particles usually do not become directly entrained by the

wind because electrostatic forces and surface crusting bind these particles together (Bagnold,

1941; Pye, 1995). However, saltating sand grains that bombard the surface can break these bonds

and eject silt- and clay-sized particles into suspension (Bagnold, 1941; Shao et al., 1993). Loess

deposits form from the settling of suspended dust particles. Loess accumulations blanket pre-

existing topography, are generally thickest proximal to their sources, and thin exponentially for

tens to hundreds of km downwind (Frazee et al., 1970). Some loess deposits exhibit dramatic

changes in local thickness, raising questions about their genesis and postdepositional history.

One recent model developed from thick Peoria Loess in the Upper Mississippi River

valley suggests that in some cases topographic traps control the thickness and distribution of

loess (Mason et al., 1999). Topographic traps such as incised river valleys promote the physical

separation of eolian sand from silt particles (Fig. 1). When the migration paths of saltating grains

encounter topographic traps, the sand becomes trapped in the valley and is restricted from further

transport downwind. Saltating sand upwind of the trap continues to eject silt and clay particles

4

which accumulate downwind of the trap. Deposition of suspension-transported dust is favored

downwind of the valley because the saltating sand has been removed from the eolian system by

the trap and is not available to re-entrain fine particles.

The topographic trap forms a stationary boundary between saltation-dominated and

suspension-dominated eolian deposits. The absence of saltating sand grains downwind of a trap

results in thicker loess that accumulates at a higher rate than loess not influenced by a trap.

Profiles of loess should record fewer incursions of coarser eolian sediment because the saltation

load has become trapped. This is in contrast to eolian systems that have a gradational boundary

between sand- and silt-dominated eolian deposits, such as Eureka Flat in south central

Washington (Sweeney, 2004). Loess being deposited downwind of eolian sand may form thick

accumulations, but these accumulations may be remobilized by saltating sand grains encroaching

from upwind. Additionally, the grain size distribution of loess downwind of a gradational

boundary may exhibit a wide range in particle sizes, reflecting the shifting sand-loess boundary.

The Palouse loess of the Pacific Northwest is the product of approximately 2 myr of

eolian reworking (Busacca, 1991) of glacial outburst flood sediments periodically deposited

during the Pleistocene (Bjornstad et al., 2001). Primary phases of loess accumulation on the

Palouse occurred during interglacial and interstadial conditions while pedogenesis dominated

during full glacial conditions (McDonald and Busacca, 1998). This pattern is in direct contrast to

many other loess systems. In the Palouse, changes in atmospheric circulation patterns during the

last glacial maximum (LGM) resulted in weakened winds and subdued eolian activity, slowing

loess deposition and thus promoting soil formation in spite of a cold climate (Sweeney et al.,

2004). Strong dust-transporting winds resumed with the recession of the ice sheets, and newly

5

emplaced glacial outburst flood sediments became available for deflation and eolian transport,

resulting in the formation of sand- and silt-rich eolian deposits.

The Horse Heaven Hills region, an elongate structural upland located along the southern

margin of the Palouse, has an arid to semi-arid climate today (~200 mm mean annual

precipitation). Local vegetation is a shrub steppe community that includes sagebrush (Artemisia

tridentata), bitterbrush (Purshia tridentata), a variety of perennial bunchgrasses (Agropyron,

Koeleria, Stipa, Poa), forbes, and invasive cheatgrass (Bromus tectorum) growing in sandy and

gravelly soils. Juniper trees (Juniperus occidentalis) grow on the north-facing canyon slope of

Juniper Canyon in sandy soils. Mollisols form in loess and Entisols form in young, sand-rich

eolian or alluvial sediments. Strong seasonal winds in the area are enhanced via topographic

funneling through the Columbia River Gorge (Gregg, 1964).

On the south, upwind-facing flank of the Horse Heaven Hills, several canyons have

incised into the underlying Columbia River Basalt that drain ephemerally into the Columbia

River. Juniper Canyon (Figs. 2 and 3), the most prominent of these canyons, is approximately

150 m deep, is about 1 km wide, and is oriented normal to the prevailing southwesterly dust-

transporting winds. The canyon serves as a geographic boundary between a continuous though

relatively thin deposit of eolian sand on the upwind side and a thick blanket of loess on the

downwind side (Fig. 2). Cold Springs Canyon (Figs. 2 and 3), located approximately 7 km south

of Juniper Canyon, is roughly 60 m deep and 500 m across. Cold Springs Canyon similarly

separates upwind eolian sand from downwind loess accumulations.

This paper evaluates the roles that Juniper and Cold Springs Canyons have played in the

distribution and thickness of loess in the southwestern Palouse. These topographic features and

their associated glacial outburst flood and eolian deposits provide an ideal test of the topographic

6

trap model of Mason et al. (1999) and in so doing provide insight into timing of loess deposition

since the LGM.

Palouse loess stratigraphy

Measurements of loess thickness and grain size across the Columbia Plateau indicate that

the prevailing winds in the Pacific Northwest have been from the southwest for at least 70,000 yr

(Fig. 4a; Busacca and McDonald, 1994). The loess is composed of numerous lithostratigraphic

units, the two most recent of which are informally named L1 (15,000 yr to present) and L2

(70,000 to 15,000 yr) (Busacca and McDonald, 1994). The L1 loess contains a modern surface

soil at its top and the Sand Hills Coulee Soil near its middle (McDonald and Busacca, 1992). The

Sand Hills Coulee Soil is a weakly formed paleosol containing a calcium carbonate horizon with

a fabric of sparse cylindrical burrows (McDonald and Busacca, 1992). The stratigraphic position

of the Sand Hills Coulee Soil within the L1 suggests that it may have formed during the late

Pleistocene or early Holocene (McDonald and Busacca, 1992). The L1 and L2 loess are

separated by the Washtucna Soil, a paleosol that formed in the upper part of the L2 loess

between 40,000 and 20,000 yr (Richardson et al., 1997). The Mt St Helens set S tephra (15,400 ±

100 cal yr B.P.; Mullineaux, 1986; CALIB v4.4, Stuiver and Reimer, 1993; calibration data set

of Stuiver et al., 1998) is sometimes preserved above the Washtucna Soil; together this tephra

and paleosol are important chronostratigraphic and pedostratigraphic markers (Busacca et al.,

1992; McDonald and Busacca, 1992). The Washtucna Soil is easily recognized by its light-

colored, resistant petrocalcic horizon that has a fabric of continuous cylindrical burrows. In many

locations it has a laminar cap cemented by calcium carbonate (McDonald and Busacca, 1990).

The cylindrical burrow fabric within the Washtucna and Sand Hills Coulee soils was generated

7

by nymphs of burrowing cicadas (Cicadidae) that fed on the perennial, woody roots of shrubs

such as sagebrush (Artemisia), the dominant vegetation type on the Columbia Plateau during the

last glacial maximum (O’Geen and Busacca, 2001; Blinnikov et al., 2002; Whitlock et al., 2000).

Sagebrush also was abundant during the development of the Sand Hills Coulee Soil (O’Geen and

Busacca, 2001; Blinnikov et al., 2002).

Glacial outburst flooding

The last episode of late Pleistocene glacial outburst floods occurred approximately

18,000 to 14,500 cal yr B.P. (15,300 to 12,700 14

C yr B.P.; Waitt, 1985) after deposition of L2

had been completed. These late glacial floods and the sediment they deposited in basin areas

triggered the onset of L1 deposition (Busacca and McDonald, 1994). Glacial outburst flooding

has occurred in the Pacific Northwest for nearly 2 myr (Bjornstad et al., 2001) and has

periodically replenished sand- and silt-rich sediment in basins that then were mobilized by the

wind (Busacca and McDonald, 1994; Sweeney et al., 2002). Major areas of slackwater sediment

deposition (fine-grained sediment deposited in backflooded areas) upstream of Wallula Gap, a

major constriction along the Columbia River (Fig. 3), include the Walla Walla Valley, Pasco

Basin, and Yakima Valley (Waitt, 1985). Once flood waters funneled through Walulla Gap, they

encountered another constriction along the Columbia River at Rowena Gap, resulting in

slackwater deposition in the Umatilla Basin of northern Oregon (O’Connor and Waitt, 1995;

Benito and O’Connor, 2003).

8

Methods

We mapped the distribution of eolian and glacial outburst-flood sediments on the Horse

Heaven Hills in proximity of the Columbia River. We made transects downwind of Juniper and

Cold Springs Canyons to measure loess thickness. The thickness of the L1 loess and other

surficial sediment was determined at over 40 sites using a hand auger. Selected sites are depicted

on Fig. 3. Sediment samples were taken from hand auger cuttings at approximately 30 cm

intervals at each site. Sites were georeferenced using a global positioning system (GPS).

In the lab we measured grain-size distributions of samples using a Malvern Mastersizer S,

a laser diffractometer that measures volume percent of particles in 64 size classes from 0.05 to

850 µm. Samples were pretreated prior to analysis with sodium acetate to dissolve soil

carbonates and with hydrogen peroxide warmed on a hot plate to oxidize organic matter.

Samples were then rinsed in de-ionized water, centrifuged, and decanted. Each sample was

dispersed with sodium hexametaphosphate and analyzed in a de-ionized water suspension with

no sonication. Samples were described based on mean grain size and sand content. Volcanic

tephras were analyzed with a Cameca Camebax electron microprobe at Washington State

University, with an acceleration voltage of 15 kV, a beam current of 10 nA, and a beam diameter

of 6 µm. Glass shards from each sample were analyzed for Si, Al, Fe, Ti, Na, K, Mg, and Ca

oxides and elemental Cl and compared to a library of Pacific Northwest tephras and standards

(Clague et al., 2003; Foit et al., 1993) using similarity coefficients (Borchardt et al., 1972) for

each sample.

9

Results

Glacial outburst flood features

We did not find glacial outburst flood features above 350 m asl, consistent with height

estimates of the latest Pleistocene paleofloods (Fig. 2; O’Connor and Baker, 1992). Erosional

features that we documented in the study area include scabland topography composed of basalt

exposures and steep scarps cut into thick loess (Fig. 2). Floods scoured loess and preexisting

sediment for a distance of more than 2 km east of the Columbia River in the Juniper Canyon

area, and more than 8 km east at lower elevations in the Cold Springs Canyon area. Floods

deposited basalt-rich gravel up to 2 km upstream from the mouth of Juniper Canyon. Subsequent

flooding and fluvial incision of this flood gravel generated a terrace that parallels the southern

slope of Juniper Canyon and lies 50 m above the canyon base. Active and stabilized dunes

mantle this terrace. Flood-transported basalt-rich sand, reworked loess, and reworked calcium

carbonate soil nodules drape the landscape upwind of the canyon.

Eolian Features

Sand-rich, gravel-bearing outburst-flood deposits have been reworked by wind to form

thin sand sheets and parabolic dunes (Fig. 3; JC1, JC24, JC28). Dunes have migrated to the

southern (upwind) margin of, and locally into, Juniper Canyon (Fig. 5a). Basalt-rich sand sheets

are vegetated or agriculturally modified and overlie outburst flood sediment. Basaltic granule

lags are exposed in zones of net deflation among the parabolic dunes and sand sheets. Above the

maximum elevation of outburst flooding, blowouts behind parabolic sand dunes have exposed

the laminar carbonate cap of the Washtucna Soil. In one locality, yardangs about 0.25 m high

have formed in the relatively resistant Washtucna Soil; these deflational features are surrounded

10

by a lag composed of carbonate-rich nodules and basalt granules (Yardang, Fig. 3; Fig. 5b).

Cliff’s Blowout is on the eastern margin of eolian sand transport and exposes eolian sand above

and below a 0.20 m thick bed of the Mazama tephra (ca. 7600 cal yr B.P.; Zdanowicz et al.,

1999) (Figs. 3 and 5c).

The southern, north-facing slope of Juniper Canyon is mantled by eolian sand that has

avalanched into numerous dry sand flows. At the base of the canyon, sand flow sediment is

exposed above and below a 0.20 m accumulation of Mazama tephra (Fig. 5d). In some places,

the sand flow sediment has accumulated on fluvial or outburst-flood terraces within the canyon

and has been remobilized into parabolic dunes. The northern, south-facing slope of the canyon is

composed mostly of basalt outcrops with a thin (<10 cm) loess cover.

Upwind of Juniper Canyon, L1 loess has accumulated over L2 loess in a number of

locations (Fig. 2). This includes L1 loess that accumulated downwind of the Cold Springs

Canyon topographic trap (Figs. 2, 6). L1 loess thins downwind from 6.4 m at JC36 north of Cold

Springs Canyon to 3.5 m at JC37 (Fig. 6). Our reconnaissance mapping confirms data from the

county soil survey that the texture of the loess fines eastward from very fine sandy loam to silt

loam (Johnson and Makinson, 1988). Sand-rich soils that include stabilized parabolic sand dunes

and areas of sand-rich loess lie upwind of Cold Springs Canyon.

Downwind of Juniper Canyon, the L1 loess is thickest where it accumulated on older

loess, and relatively thin on flood-scoured surfaces. Outburst floods generated flat-topped basalt

buttes along the margin of Walulla Gap, including at the mouth of Juniper Canyon, where pre-

existing loess and other sediments were stripped. These buttes have only thin (10’s of cm)

veneers of L1 loess (JC15; Fig. 3 and 7). In contrast, thick L1 loess accumulated and was

preserved on top of older loess where it escaped removal by outburst floods. Three thick-loess

11

sites (JC17, JC18, JC27) record up to 8 m of L1 loess (Figs. 3 and 7). All three sites are

immediately downwind of Juniper Canyon and across from areas of active eolian sand transport.

Two sites (JC34, JC39; Figs. 3 and 6) that are approximately 6 km downwind of Juniper Canyon

on the crest of the Horse Heaven Hills record about 4 m of L1 loess. Site JC38 is downwind of

the canyon but east of where sand dunes are directly upwind; it has approximately 4.2 m of L1

loess. The geographic distribution of loess sites demonstrates that the thickest L1 loess

accumulated immediately downwind of topographic traps formed by Juniper Canyon and Cold

Springs Canyon and thins dramatically within a few km (Fig. 6).

We found a paleosol at a depth of approximately 1.8 m within the L1 loess at sites JC17,

JC18, and JC36. This paleosol is characterized by a 0.6-0.8 m thick, relatively resistant zone of

calcium carbonate enrichment that contains abundant filamentous carbonate associated with root

pores and rare cylindrical peds. We found no other resistant zones or paleosols within the L1.

The L1 loess at sites downwind of the topographic traps ranges in mean grain size from

about 45 to 70 µm; it averages approximately 30% sand and coarsens upward (Sweeney, 2004).

Though poorly sorted, this loess is better sorted than most other proximal-to-source loess on the

Palouse (Fig. 8). Sites downwind of the topographic traps or within loess fields do not contain

particles coarser than 850 µm. We found sharp, 5 to 10 µm increases in mean grain size caused

by increases in very fine sand content in loess at all sites at similar depths (Sweeney, 2004). The

most notable increases occur between 2.4 and 2.7 m at sites JC36, JC37, JC38, and JC39,

between 1.8 to 2.0 m at sites JC17, JC18, JC27, and JC34, between 3.7 and 3.9 m at sites JC 37,

JC38, and JC39, and at 6 m at sites JC17 and JC27.

12

Tephra Identification

We found tephras derived from volcanoes in the Cascade Range in the loess of the Horse

Heaven Hills. These include the Mt St Helens (MSH) set S (So and Sg) and one Glacier Peak

tephra. The MSH set S tephras occur at the base of the L1 at sites JC18, JC27, and JC36 at

depths of 6.8, 7.8, and 6.4 m, respectively (Table 1). Tephras at sites JC18 and JC27 were

correlated to MSH layer Sg (similarity coefficients (SC) 0.98 and 0.97, respectively) using

standards for both So and Sg (Clague et al., 2003). The tephra at site JC36 matches So and Sg

equally well (SC for both = 0.96) and may represent mixing of tephras. The eruptions that

generated the chemically distinct So and Sg tephras likely occurred within a few decades of each

other (Clague et al., 2003) and have an age of approximately 15,400 ± 100 cal yr B.P.

(Mullineaux, 1986; CALIB 4.4, Stuiver and Reimer, 1993; calibration data set of Stuiver et al.,

1998). The age of MSH set S is confirmed by luminescence ages on loess of 14,000 yr above and

17,200 yr below the tephra at a different locality (Richardson et al., 1997).

The Glacier Peak tephra was identified at 2.4 m depth at JC36 based on a SC of 0.98 to a

Glacier Peak standard (Table 1). The Glacier Peak tephra has an age of 11,200 14

C yr B.P. (Foit

et al., 1993), with a calibrated age of 13,100 ± 200 cal yr B.P. (CALIB 4.4, Stuiver and Reimer,

1993; calibration data set of Stuiver et al., 1998). The Glacier Peak tephra has been documented

in loess-derived colluvium from the Horse Heaven Hills (A. Palmer, personal communication,

2004) as well as within the L1 loess at other sites (Busacca et al., 1992; Kemp et al., 1998).

A high-silica tephra of unknown origin was identified at 2.1 m depth in loess from JC27

(JC27-210; Table 1). A diffuse concentration of volcanic glass was found by scanning smear

slides of loess. This unknown tephra sits at a similar stratigraphic position as the Glacier Peak

tephra found in JC36, but is not a geochemical match. Correlation of tephras from site to site is

13

made difficult by erosion, changes in the efficiency of trapping vegetation, and bioturbation that

result in spatially heterogeneous preservation of tephra.

Discussion

Influences of topographic traps on loess accumulation

The combination of an abundant supply of erodible flood sediment in the Umatilla Basin,

strong prevailing southwesterly winds, and topographic traps have generated the thickest known

accumulations of L1 loess (6 to 8 m thick) from any part of the Columbia Plateau eolian system.

Where sand dunes intersect Juniper and Cold Springs canyons and the saltating sands are

trapped, thick loess has accumulated immediately downwind (sites JC17, JC18, JC27, and JC36).

Prior to this work, the greatest known thickness of L1 loess was 4.5 m at CLY-1 (Fig 4a;

Busacca and McDonald, 1994) adjacent to Eureka Flat. Topographic trapping of eolian sand may

also have influenced the thickness of loess at CLY-1. This site is separated from upwind sand

dunes by Winnett Canyon, an incised tributary of the Touchet River (Sweeney, 2004).

The thick loess that mantles the Horse Heaven Hills in the Juniper and Cold Springs

Canyon areas is proximal to its Umatilla Basin source and thins downwind, a typical pattern seen

in other loess areas such as the midwestern U.S. and Great Plains (Frazee et al., 1970; Mason,

2001) and including the Palouse (Busacca and McDonald, 1994). To test that topographic

trapping accentuated the thickness of proximal loess downwind of Juniper and Cold Springs

Canyons, we examined loess thickness at nearby sites where topographic trapping of eolian sand

was unlikely (Fig. 7). At the eastern margin of Juniper Canyon (Fig. 2), sand dunes do not

migrate up to the canyon edge. Instead, thick loess mantles the landscape on both the north and

south sides of the canyon. Thickness of L1 loess on ridge tops north and south of the canyon

14

averages approximately 4 m, in contrast to the 8 m found on ridge tops downwind of the trap.

East of the major zone of eolian sand transport, suspension fallout dominates eolian deposition

and segregation of particles by topographic trapping does not play a role in the thickness of

loess.

West of Walulla Gap in an area of the Horse Heaven Hills where stream valleys have

shallow slopes and are oriented obliquely to prevailing winds, topographic trapping of eolian

sand is minor. L1 loess is generally <3 m thick (HHH1, 2, 4, Owens1; Fig. 3). The contrast in

loess thickness demonstrates the effectiveness of topographic traps at inducing accumulation of

thick loess downwind of the canyon. Loess downwind of Juniper Canyon is nearly two times as

thick as areas with no topographic trap. The 4 m of L1 loess that has accumulated in places near

Juniper Canyon in the absence of a topographic trap suggests that Umatilla Basin has been a

major producer of wind-blown silt.

The segregation of eolian sand from loess by a formidable topographic trap enhances the

sorting of loess. Thick loess downwind of the canyons is better sorted than loess accumulating

downwind of its source in areas where there is no topographic trap. On average, grain size

distribution of loess downwind of Juniper and Cold Springs Canyons has a standard deviation

(sorting) of 1.5 Φ, compared to 2.0 Φ for loess proximal to Eureka Flat in south-central

Washington where traps do not occur (Fig. 8). The proximity of dunes to loess combined with

strong winds can result in coarser grained loess (Porter and An, 1995; Vandenberghe and

Nugteren, 2001). Loess that has accumulated downwind of dunes on Eureka Flat has a mean

grain size that varies by >50 µm throughout its thickness. The mean grain size of loess

downwind of Juniper Canyon varies by < 25 µm (Fig. 8). The spikes of very fine sand in the

loess downwind of the topographic traps may represent brief periods of strong winds where the

15

very fine sand was transported temporarily by suspension. The coarsening-upward trend within

this loess may reflect the gradual encroachment of eolian sand to the upwind margin of the

canyons since about 15,000 yr ago.

Influence of glacial outburst floods and paleoclimate on loess formation

The preservation of thick loess downwind of topographic traps is aided by dust-trapping

vegetation. Paleoecologic studies of the Columbia Plateau suggest that the region supported

bunchgrass steppe or sagebrush steppe for the last 100,000 yr (Blinnikov et al., 2002; Whitlock

and Bartlein, 1997; Whitlock et al., 2000). Glacial outburst flooding at Juniper Canyon modified

the landscape that influenced where loess could accumulate, mostly by removing soil and

preexisting plant cover. Areas of vegetated loess that escaped erosion from flooding near Juniper

Canyon allowed continued loess accumulation whenever dust was being generated, resulting in a

thick and continuous record. Below the maximum elevation of outburst floods, much of the pre-

existing loess was stripped, leaving either exposed basalt or outburst-flood sediment. Where

gravelly to sandy flood sediment was deposited, loess was not able to accumulate until

vegetation re-colonized the area, generating a time lag between flooding and the onset of loess

accumulation. On bedrock exposures with extremely sparse vegetation, little to no loess has

accumulated in 15,000 yr. These surfaces are characterized by sparse vegetation that roots in

fractures within the basalt bedrock, providing little plant canopy to protect dust that falls on the

surfaces from being removed by rain splash, runoff, and re-entrainment by wind. Areas with

saltating sand additionally re-entrain dust particles, resulting in loess accumulation and

preservation only beyond areas of active saltation.

16

When dust accumulation rates decrease, longer residence time of soil development

processes in any volume of loess is reflected by more strongly developed soils. We interpret the

paleosol in L1 at sites JC17, JC18, and JC36 as the Sand Hills Coulee Soil based on its

stratigraphic position and features that are consistent with its type locality elsewhere on the

Columbia Plateau (McDonald and Busacca, 1990). The Glacier Peak tephra is near the base of

the buried Bk horizon of the Sand Hills Coulee Soil at site JC36. Pedogenic features such as

cylindrical nodules and calcium carbonate cement, although stratigraphically at the same interval

as the tephra, post-date the tephra, suggesting decreased dust accumulation and onset of soil

development after 13,100 cal yr B.P. Cylindrical nodules formed by burrowing cicadas suggest

sagebrush steppe vegetation at the time of soil formation (O’Geen and Busacca, 2001). Soils on

the Columbia Plateau that include cicada burrows have been linked to dry conditions (O’Geen

and Busacca, 2001).

The Younger Dryas Stade (YD; ca. 13,000 to 11,500 cal yr B.P.; Whitlock and Grigg,

1999) could have provided ideal conditions for the expansion of sagebrush steppe on the

Columbia Plateau and these may have been recorded in the Sand Hills Coulee Soil. The position

of the Glacier Peak tephra, however, may be a poor estimate for timing of soil formation,

considering the time required to accumulate loess above the tephra and then to form the soil. The

Sand Hills Coulee Soil alternatively could be associated with warm and dry conditions of the

early to middle Holocene that have been recorded by pollen extracted from Carp Lake sediments

on the western margin of the Columbia Plateau (Whitlock, et al., 2000). The degree of

development of the Sand Hills Coulee Soil, including burrow density and carbonate content,

apparently represents only a few thousand years of pedogenesis (McDonald and Busacca, 1992),

consistent with either of the above climate intervals.

17

Heightened dune activity during the early to middle Holocene is indicated by the

presence of Mazama tephra in eolian sand of Cliff’s Blowout and within dry sand flow deposits

in Juniper Canyon. Dune activity may correlate to loess accumulation downwind of Juniper

Canyon. Early to middle Holocene eolian activity has also been documented in dunes of the

central Columbia Plateau (Gaylord et al., 2001).

Mass accumulation rates of loess

Mass accumulation rates (MARs) from loess in the Horse Heaven Hills region indicate

that the bulk of L1 loess accumulated during the latest Pleistocene. The MAR is calculated by

multiplying the rate of dust accumulation (m yr-1

) by the bulk density of the loess (Bettis et al.,

2003), in this case 1.40 g cm-3

on average. MSH set S tephra (15,400 cal yr B.P.) and the Glacier

Peak tephra (13,100 cal yr B.P.) bracket 4.0 m of loess at JC36 downwind of Cold Springs

Canyon (Figs. 2, 5), yielding an average MAR of about 2500 g m-2

yr-1

. L1 loess downwind from

Juniper Canyon (Figs. 2, 6) contains the MSH set S tephra at its base and is thicker than at site

JC36, thus it likely has a higher MAR, but it does not contain Glacier Peak tephra. MARs for L1

loess are high when compared to most late glacial loess worldwide, (see Roberts et al., 2003), but

are less than MARs of ~11,500 g m-2

yr-1

between 18,000 and 14,000 yr calculated for proximal

late-glacial loess in Nebraska (Roberts et al., 2003). The timing of peak accumulation rates in the

Palouse lag behind those on the Great Plains by a few thousand years.

The MAR for L1 loess that accumulated above the Glacier Peak tephra at JC36 is

approximately 250 g m-2

yr-1

, an order of magnitude lower than the MAR below the tephra. The

MAR for loess above the Glacier Peak tephra is an average value that assumes a gradual

18

accumulation rate of 2.4 m in 13,100 yr. Additional dating of closely spaced intervals of loess

above and below the Glacier Peak tephra is necessary to calculate more accurate MARs.

The timing of increased dust production (15,400-13,100 yr) corresponds to the return of

strong prevailing westerly flow for the Pacific Northwest following a phase of weakened

westerly flow caused by the glacial anticyclone during the LGM (Bartlein et al., 1998). The

weakened prevailing winds caused by the anticyclone decreased dust production on the

Columbia Plateau (Sweeney et al., 2004). Replenished source sediment emplaced by glacial-

outburst flooding became available for deflation by re-invigorated westerly winds after the

anticyclone dissipated. Figure 9 is a summary diagram that links the timing of loess

accumulation, soil formation, and associated paleoclimates interpreted from this study.

Umatilla Basin influence on the Palouse loess

The Umatilla Basin is a major depocenter of sand- and silt-rich glacial outburst flood

sediments (O’Connor and Waitt, 1995; Benito and O’Connor, 2003), but its role as a source of

the Palouse loess has not been evaluated until now. At least 8 loess units are preserved at the

Helix site east of Juniper Canyon (Tate, 1998). A TL age of 158,000 +/-16,800 yr was obtained

near the top of the sixth unit below the surface (Richardson et al., 1997) suggesting that the

Umatilla Basin has a long record of producing dust, rivaling other source basins in the formation

of the Palouse.

The discovery of 4 m of older L2 loess preserved upwind of Juniper Canyon suggests that

topographic trapping may not have influenced loess thickness in this area for at least part of the

time of L2 loess deposition, approximately 70,000 to 15,000 yr. Climatic conditions during L2

time were arid to semi-arid, as inferred from thick accumulations of L2 loess across the

19

Columbia Plateau (Busacca and McDonald, 1994). There is no stratigraphic evidence that eolian

sand migrated up to Juniper Canyon during L2 time, suggesting that dust-producing saltating

sand occurred farther upwind in the Umatilla Basin. Such reduced sand transport is attributed to

denser plant cover, higher soil moisture, or some other physical restriction that resulted in L2

loess accumulating both upwind and downwind of Juniper Canyon. Late Pleistocene outburst

flooding that scoured L2 loess from the western end of Juniper Canyon also deposited silt- and

sand-rich flood sediment. The distribution of flood sediment has contributed to dune formation

immediately upwind of the canyon, and has allowed the topographic trap to function for greater

than 15,000 yr.

Regional loess thickness trends suggest that the Umatilla Basin has influenced the

evolution of the Palouse loess downwind of the Horse Heaven Hills. Thin loess (<1 m) that

mantles glacial outburst flood sediments in the Walla Walla Valley (WWV) north of the Horse

Heaven Hills is derived from the Umatilla Basin (Fig. 4a). Loess derived from Eureka Flat (EF)

also conceivably had a complementary source from the Umatilla Basin. Figure 4b depicts

changes in L1 loess thickness with distance downwind from both the Umatilla Basin and Eureka

Flat. The exponential downwind thinning of loess derived from the Umatilla Basin continues

downwind of the Walla Walla Valley into loess derived from Eureka Flat, suggesting that thick

loess sites such as CLY-1 likely contain <1 m of loess derived from the Umatilla Basin (Fig. 4b).

The superimposed plumes of dust depicted in Fig. 4a, whose sources are separated by the Horse

Heaven Hills, combined to contribute to the accumulation of the Palouse loess.

20

Conclusions

Glacial outburst flooding during the late Pleistocene produced the main sources of

sediment for the Palouse loess. The floods emplaced extensive sand- and silt-rich source

sediments upwind of the Horse Heaven Hills in the Umatilla Basin. The combination of dry and

windy conditions and the presence of topographic traps for eolian sand resulted in high MARs of

dust for a short interval during the late glacial between about 15,400 and 13,100 cal yr B.P. High

rates of dust production were followed by a short interval of landscape stability and soil

formation represented by the Sand Hills Coulee Soil. The formation of this soil may be linked to

drier conditions in the latest Pleistocene or early to middle Holocene. MARs from 13,100 cal yr

B.P. to present are an order of magnitude lower than those in the latest Pleistocene.

Juniper and Cold Springs Canyons, deep canyons oriented perpendicular to the prevailing

wind, acted as topographic traps for migrating sand dunes, segregating saltating particles from

the eolian system and allowing late Pleistocene to Holocene L1 loess to reach thicknesses of 6 to

8 m downwind. As such, the depositional and pedologic record of loess downwind of the

canyons provide a positive test of the topographic trap model (Mason et al., 1999) indicating that

it is applicable beyond the Upper Mississippi valley. Results from this study expand the scope of

the model’s application to explain relatively well-sorted loess downwind of traps and

demonstrate how trapping, in conjunction with dry and windy climates, increases accumulation

rates of loess.

Acknowledgements

We thank Joe Mason, Bill Zanner, and Alan Gillespie for constructive comments. Ed

Brook reviewed and earlier draft of this manuscript. Thanks to Brandt Halver and Luke Lemond

21

for field assistance, Denise Honn and Rick Trotman for laboratory assistance, and Nick Foit for

tephra analysis and helpful discussion. Thanks to Phil and Mike Hawman, Cliff Bracher, and

Tom Peterson for allowing access to their properties. This research was supported by grants from

the National Science Foundation (ATM-0214508 to Busacca and Gaylord), the Geological

Society of America, and the Columbia Plateau Wind Erosion Project (USDA).

22

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Figure Captions

Fig. 1. Topographic trap model for thick loess accumulation. Eolian sand is trapped by an incised

stream valley, allowing thick loess to accumulate from suspension fallout downwind of the trap.

After Mason et al. (1999).

Fig. 2. Shaded relief map of the Juniper Canyon area, northern Oregon, depicting the relation of

selected site locations to topographic traps and glacial outburst flooding features. Areas east of

the dashed line depict elevations above glacial outburst flood levels in the region (from data

collected by O’Connor and Baker (1992)). The stippled area depicts the zone of active eolian

sand saltation transport. Inset map shows location of Juniper Canyon (star), major constrictions

along the Columbia River, and the location of Umatilla Basin (UB).

Fig. 3. USGS aerial photograph showing all sample locations and major geographic features.

Fig. 4. A. Isopach map of the regional L1 loess thickness, showing a double plume of dust

sourced from the Umatilla Basin in Oregon and basins in south-central Washington. Sample sites

CLY-1 and JC27 from figure B are included. The dashed line refers to the loess thickness profile

illustrated in figure B. S = Spokane, P = Pullman, EF = Eureka Flat, WWV = Walla Walla

Valley. Modified from Busacca and McDonald (1994). B. L1 loess thickness versus distance.

Diamonds depict loess thickness downwind of the Umatilla Basin (UB); squares depict loess

thickness downwind from Eureka Flat (EF) derived from section line in 4A. Loess thickness

trends indicate that the UB likely contributed loess to areas downwind of Eureka Flat (dotted

line).

31

Fig. 5. A. Sand dunes trapped by Juniper Canyon, looking east. Prevailing wind directions are

from right to left. Arrow depicts direction of dune migration. Juniper trees are approximately 4-5

m tall. B. Small yardang formed in the Washtucna Soil. Area is active zone of deflation upwind

of Juniper Canyon. Field book is 19 cm high. C. Cliff’s Blowout south of Juniper Canyon,

revealing exposures of Mazama tephra within eolian sand. D. Dry sand flows along the north-

facing slope of Juniper Canyon. Mazama tephra is exposed between two older dry sand flows

deposits. Note 1.0 m shovel for scale.

Fig. 6. Transects showing downwind thinning of loess from topographic traps. Distances labeled

between sites. Mt St Helens set S and Glacier Peak tephras were used to correlate between sites

and calculate mass accumulation rates. Refer to Figure 3 for site locations.

Fig. 7. Transect demonstrating the efficiency of Juniper Canyon as a topographic trap that

generates very thick loess. Sites HH1, JC37, and JC38 have thinner loess not affected by

topographic trapping. Refer to figure 3 for site locations.

Fig. 8. Mean grain size (φ) plotted against standard deviation (sorting, φ) for proximal loess from

Juniper Canyon (JC) and Eureka Flat (EF), south-central Washington.

Fig. 9. Summary diagram relating the time frame of loess accumulation to regional stratigraphy,

soils, tephras, and paleoclimate in the Horse Heaven Hills area based on this work, Busacca et

al., 1992, and McDonald and Busacca, 1992.

32

Table 1. Glass chemistry of tephras from loess.

Table 1

Glass chemistry of tephras from loess

Oxide JC27-240 JC27-780 JC18-680 JC36-210 JC36-640 MSH So

stda

MSH Sg std

a

Glacier Peak std

b

SiO2 79.80 (0.51)c 76.48 (0.26) 76.55 (0.22) 77.43 (0.23) 76.63 (0.46) 77.08 (0.28) 76.50 (0.14) 77.47 (0.38)

Al2O3 12.93 (0.19) 13.87 (0.10) 13.62 (0.15) 12.51 (0.10) 13.49 (0.30) 13.39 (0.18) 13.80 (0.09) 12.59 (0.16)

Fe2O3 1.08 (0.06) 1.32 (0.10) 1.28 (0.04) 1.20 (0.08) 1.25 (0.09) 1.22 (0.04) 1.29 (0.03) 1.30 (0.13)

TiO2 0.21 (0.02) 0.16 (0.02) 0.16 (0.02) 0.22 (0.02) 0.15 (0.03) 0.17 (0.02) 0.16 (0.02) 0.21 (0.03)

Na2O 1.91 (0.22) 3.89 (0.15) 3.91 (0.12) 3.64 (0.11) 4.07 (0.15) 4.08 (0.17) 4.11 (0.12) 3.56 (0.15)

K2O 2.41 (0.17) 2.22 (0.05) 2.34 (0.10) 3.24 (0.11) 2.45 (0.40) 2.24 (0.07) 2.14 (0.04) 3.21 (0.28)

MgO 0.24 (0.02) 0.31 (0.02) 0.34 (0.03) 0.27 (0.03) 0.30 (0.07) 0.28 (0.02) 0.32 (0.02) 0.26 (0.05)

CaO 1.24 (0.07) 1.62 (0.06) 1.66 (0.05) 1.28 (0.04) 1.53 (0.27) 1.44 (0.06) 1.59 (0.02) 1.26 (0.15)

Cl 0.18 (0.03) 0.10 (0.01) 0.11 (0.02) 0.21 (0.02) 0.12 (0.02) 0.10 (0.02) 0.09 (0.01) 0.16 (0.04)

TOTALd 100 100 100 100 100 100 100 100

Number of shards analyzed

12 18 19 16 17 22 20 18

Tephra ID ??? MSH Sg MSH Sg Glacier Peak MSH S

aMt St Helens set So and Sg standards, Clague et al. (2003)

bGlacier Peak standard, Foit et al., (1993), Wild Cat Lake, WA

c Standard deviations of the analyses are given in parentheses

dAnalyses normalized to 100 weight percent

33

Fig 1

34

Fig 2

35

Fig 3

36

Fig 4

37

Fig 5

38

Fig 6

39

Fig 7

40

Fig 8

41

Fig 9