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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, [email protected]
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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.
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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