Rates and processes of soil development on Quaternary terraces
in Cajon Pass, California
LESLIE D. McFADDEN Department of Geology, University of New Mexico, Albuquerque, New Mexico 87131
RAY J. WELDON II U.S. Geological Survey Branch of Engineering Seismology and Geology, M.S. 977, 345 Middlefield Road, Menlo Park,
California 94025
ABSTRACT
Field and laboratory analyses of soils on 11
well-dated fluvisil terraces spanning the past
0.5 m.y. demonstrate that a threshold governs
changes in several morphological and chemical characteristics of increasingly older soils.
Correlations with respect to time among iron
species, soil morphology, and soil silt and clay
demonstrate that the chronosequence at
Cajon Pass reflects primarily an evolutionary,
largely time-dependent trend and does not reflect differences in external factors such as
climate. Most of the soil development on
Holocene terraces of the Cajon Pass area is
due to physical incorporation of eolian dust
and organic material into initially very permeable gravels. This process decreases soil
permeability and is conducive for an increase
in the magnitude of chemical weathering.
Latest Pleistocene and older Pleistocene soils
have developed clay and authigenic iron oxide-rich B horizons at the expense of organicmatter-rich A horizons and color B horizons
as the extent of chemical weathering has in-
creased. This conversion of the soil from a
noncolloidal system to a much more colloidal
system takes place over a relatively short period of time (<4,000 yr) and is herein defined
as a type of pedologic threshold. In the Cajon
Pass area, the attainment of the threshold and
subsequent development of the argillic B horizon of soils on latest Pleistocene terraces
occurred during the Holocene; thus, the absence of argillic horizons in soils on Holocene
terraces is attributable to simply their
younger age rather than to the Pleistoceneto-Holocene climatic change. The threshold is
a function of several variables, including influx rate of eolian dust and initial soil permeability; therefore, the time required to
attain the threshold will vary in chronosequences characterized by geomorphic or
geographic settings that are different from
conditions found in Cajon Pass.
INTRODUCTION
The disciplines of tectonic geomorphology,
neotectonics, and paleoclimatology often rely on
35
soil development to date young deposits because
radiometrically datable materials or index fossils
are absent or scarce in most terrigenous Quaternary deposits and soils occur commonly on
Quaternary deposits. New techniques such as
uranium-trend and thermoluminescence dating
are experimental, hence uncertain. The geochronological information provided by soil
development is especially appropriate in cases in
which the distinction between the ages of a
geomorphic surface and its substrate is a critical
consideration.
Soils are highly complex natural systems and
are affected by variables that include topography, parent materials, vegetation, climate,
and time. The state factor approach of Jenny
(1941) provided a sound basis for using pedologic data to infer age by holding the influence
of the other factors constant. Many studies have
shown that certain soil properties are related to
soil age; among these properties are morphology
of calcic horizons (Gile and others, 1966), mass
of secondary carbonate (Machette, 1978,1985),
thickness of stone weathering rinds (Colman
and Pierce, 1981), soil morphological properties
such its horizon thickness and clay content
(Bockheim, 1980), total soil morphology
(Harden, 1982; Birkeland, 1984), and iron oxide
content (McFadden and Hendricks, 1985). Absolute rates of soil development, however, have
not been determined in most soil chronosequence studies because few absolute ages of the
soils are known. Some variables, such as climate
or influx of eolian dust, have changed with time
and may have significantly affected the rate of
^
Figure 1. Map showing location and general geologic setting of
the Cajon Pass study area (cp) and selected geographic features referred to in this paper. Los Angeles (LA), San Bernardino (SB), Salton
Sea (SS), San Andreas faulli (SAF), and Cleghorn fault (cf) are
included for reference.
Additional material for this article (tables) may be obtained free of charge by requesting Supplementary Data 87-07 from the GSA
Documents Secretary.
Geological Society of America Bulletin, v. 98, p. 280-293,9 figs., 3 tables, March 1987.
280
SOIL DEVELOPMENT ON QUATERNARY TERRACES, CALIFORNIA
soil development (Bockheim, 1980). For example, rates of clay or carbonate accumulation may
in part reflect changes in the rate and magnitude
of dust influx (Gile and others, 1966,1981; Yaalon and Ganor, 1973; Colman, 1982; McFadden and Tinsley, 1982, 1985; Machette, 1985)
as well as the intensity of leaching and associated chemical weathering, all of which strongly depend on climate (Rogers, 1980; Jenny,
1980; Birkeland, 1984). For these reasons, soil
chronosequence data do provide broad age estimates but are most useful for determining the
relative ages of geomorphic surfaces.
Recent studies of the late Cenozoic history of
the Cajon Creek area in the Transverse Ranges,
California (Fig. 1), provide absolute ages of
Holocene and Pleistocene deposits and geomorphic surfaces (Weldon, 1986; Weldon and Sieh,
1985). We have described and analyzed soils on
11 well-dated fluvial terraces in this area that
range in age from 47 to -500,000 yr (Fig. 2).
The climate of this region is classically Mediterranean, characterized by hot, dry summers and
cool, moist winters; annual precipitation generally ranges from 630 to 730 mm (Alhborn,
1982). Climatic variation across the study area is
minimal. With one exception, elevations of terraces in the study area range from 710 to 950 m
above sea level. The oldest and best described
soil in the study area is best preserved at an
elevation of 1,220 m, where the annual precipitation is only 430 mm. Terrace sediments are
composed primarily of albite-epidote-micachlorite schist and melanocratic to leucocratic
granitic rocks. Parent materials, vegetation, and
topographic relief are similar on most of these
surfaces, which gives us an opportunity to determine the rates at which soil properties have
281
developed during the past 0.5 m.y. and to evaluate how soil-forming variables such as climatic
change and eolian influx have influenced rates of
soil development during the late Quaternary.
QUATERNARY HISTORY OF THE
CAJON CREEK AREA
Cajon Creek is in the central Transverse
Ranges and has formed a flight of terraces across
the San Andreas fault (Figs. 1 and 2). To characterize the tectonic deformation associated with
the San Andreas fault, the Quaternary deposits
in the Cajon Creek drainage were mapped in
detail and dated using 14C, magnetic stratigraphy, and fossils (Weldon and Sieh, 1985;
Weldon, 1986). The ages of the surfaces that
formed on most late Quaternary deposits, given
in Table 1, can be closely constrained using 18
2000-1
Figure 2. Schematic section through Cajon Creek area, showing ages of deposits, surfaces, and the height of surfaces above the active
channel of Cajon Creek. Details of terrace deposits (Qoa) and surfaces (Qt) are tabulated in Table 1 and discussed in the text.
TABLE 1. AGE, TEXTURAL, AND CHEMICAL CHARACTERISTICS OF KEY QUATERNARY SOILS IN THE CAJON PASS AREA, SOUTHERN CALIFORNIA
Particle size, <2 mm (%)
Deposit,
terrace
Surface age
(yr B.P.)
Number
Profile
horizon
Depth
(cm)
Color*
(dry; moist)
Sand
Silt
Clay
pH
6.3
7.0
Qal-0
47
RW-9
A
Cu
0-27
27+
2.5Y 5/2; 4 / 2
2.5Y 6/2; 4 / 2
90.4
93.5
9.6
6.1
tr
0.4
Qoa-a,
Qt-6
275
+385
-75
RW-18
O
Al
2A2
2AC
2Cox
2Cu
1.0-0
0-3
3-18
18-34.0
34-55
55+
2.5Y 4/2; 5Y 2.5/1
2.5Y 5/2; 5Y 3/1
2.5Y 5/2; 5Y 3 / 2
2.5Y 6/4; 3 / 2
5Y 6/2; 5Y 3 / 2
72.3
82.0
83.7
89.4
91.2
23.7
15.2
13.7
7.7
6.9
4.0
2.8
2.6
2.9
1.9
4.6
4.9
5.4
5.6
5.7
Qoa-a,
Qt-6
275
+385
-75
RW-10
Al
A2
AC
Cox
Cu
0-5
5-21
21-26
26-33
33+
10YR
10YR
10YR
10YR
10YR
95.1
84.4
88.1
96.0
91.5
4.9
12.8
9.6
3.4
6.9
0.5
2.8
2.3
0.6
1.6
4.9
5.4
5.5
5.7
6.4
Qoa-c,
Qt-5
5900
± 900
RW-12
Al
A2
Bw
Coxl
Cox2
Cu/Cox
0-18
18-35
35-63
63-73
73-90
90+
10YR-2.5Y 5/2; 2 / 2
10YR-2.5Y 5/2; 3 / 2
2.5Y 6/3; 5 / 4
2.5Y 6/2; 5 / 4
2.5Y 6/2; 4 / 2
5Y 5/2; 4 / 2
84.5
87.1
91.5
88.5
89.8
94.8
13.3
10.9
7.6
10.0
8.8
3.9
2.2
2.0
0.9
1.5
1.4
1.3
5.1
5.0
5.2
5.4
5.7
5.8
Qoa-c,
Qt-4
7150
± 1200
RW-15
Al
A2
0-6
6-25
69.0
68.8
30.0
31.2
1.0
tr
5.2
5.6
BA
25-32
2.5Y 4/2; 3 / 2
10YR 4/3;
2.5Y/10YR 3 / 2
10YR 5/3;
2.5Y/10YR 3 / 3
10YR 5/4;
2.5Y/10YR 3 / 2
2.5Y 5/4;
2.5Y/5Y 4 / 2
5Y 6/2; 4 / 2
79.7
20.3
tr
5.6
86.1
13.9
tr
5.8
89.3
10.7
tr
5.9
90.4
9.6
tr
6.2
10YR 5/2; 2 / 2
10YR 5/3; 2 / 2
10YR 5/3; 4 / 3
10YR 5/4; 3 / 4
10YR-2.5Y 6/4;
10YR 4 / 3
10YR-2.5Y 4/4;
2.5 Y 4 / 4
2.5Y 6 / 3
5Y 5/2; 3 / 2
2.5Y 4 / 2
70.8
68.2
71.6
71.3
82.5
28.8
30.8
27.4
26.8
16.5
0.5
1.0
1.0
1.8
1.0
4.6
4.9
5.0
5.1
5.2
87.2
12.8
tr
5.2
94.9
91.0
5.1
9.0
tr
tr
5.5
5.6
10YR 5/3; 3 / 2
10YR 5/3; 3 / 3
8.75YR 5/4; 3 / 4
10YR 5/6; 4 / 3
2.5Y 5/2; 3 / 2
5Y 5/2; 3 / 2
70.0
69.6
69.0
74.0
94.8
94.8
27.2
26.8
25.6
20.1
4.2
4.7
2.8
3.6
5.4
5.9
1.0
0.5
5.4
5.7
5.4
5.3
5.4
5.3
10YR 4/3; 3 / 2
10YR 5/3; 3 / 3
10YR 5/4; 3 / 3
10YR 5/3; 4 / 3
8.75YR 5/4;
8.75YR 4 / 4
10YR 5/4; 10YR 4 / 4
10YR-2Y 5/4; 10YR 4 / 3
2.5Y 5/4; 2.5Y 4 / 2
2.5Y 5/4-5Y 5/3;
2.5Y 4 / 2
65.3
69.6
71.4
71.0
74.6
29.6
25.5
22.4
19.6
15.1
5.1
4.9
6.2
9.4
10.3
5.1
5.1
5.3
5.3
5.4
75.5
80.7
84.4
85.4
14.8
11.7
9.3
9.2
9.7
7.6
6.3
5.4
5.3
5.4
5.2
5.3
Qoa-c,
Qt-3
8350
+900
-500
RW-13
Bw
32-45
Cox
45-89
Cu
89+
Al
A2
AB
2Bwl
2Bw2
0-2.5
2.5-7.5
7.5-15
15-26
26-37
2BC
37-53
2Cox
2Cu
53-100+
3,000+
4/3;
5/3;
5/3;
6/4;
6/3;
3/2
2/2
4/3
4/4
5/4
Qoa-c
Qt-2
11.500
+2000
-3000
RW-6
Al
A2
2Btl
2Bt2
2Coxl
2Cox2
0-6
6-24
24-48
48-59
59-85
85+
Qoa-c,
Qt-1
12,400
± 1000
RW-17
O
Al
A2
Btl
Bt2
Bt3
I M
0-3.5
3.5-12
12-21
21-37
37-50
Bt4
BC
Coxl
Cox2
50-65
65-79
79-99
99-110+
Ol
BA
2Btl
2Bt2
2Bt3
2BC
2Coxl
2Cox2
7-0
0-9
9-42
42-77
77-99
99-140
140-190
190+
7.5YR 5/4; 3 / 4
5YR 4/6; 4 / 4
5YR 5/6; 4 / 4
6.25YR 5/4; 4 / 4
7.5YR 6/6; 4 / 6
8.75YR 6/4; 4 / 4
10YR 7/4; 4 / 4
60.0
29.0
52.5
60.0
81.3
88.0
93.5
22.7
46.8
25.7
21.6
16.9
12.0
6.4
17.3
24.2
21.8
18.4
1.8
tr
0.1
5.5
5.9
5.8
5.7
5.5
5.3
5.3
2BAt
2Btl
2Bt2
2Bt3
2Bt4
2Bt5
2Bt6
2Bt7
2Cox
0-13
13-36
36-54
54-142
142-183
183-335
335-701
701-1,460
1,460+
2.5YR 4/6; 4 / 6
2.5YR 4/6; 4 / 4
3.75YR 4/4; 5YR 5 / 6
5YR 5/6; 4 / 6
5YR 5/6; 4 / 6
5YR 5/6; 4 / 6
5YR 5/6; 4 / 6
6.25YR 5/6; 4 / 6
10YR 5/6; 4 / 6
56.2
47.2
62.7
67.4
75.4
72.1
77.7
15.5
18.3
15.0
17.0
10.8
12.6
10.3
28.3
34.5
22.3
15.6
13.8
15.3
12.0
93.8
4.7
1.6
4.4
4.3
5.5
5.8
5.3
5.4
5.2
5.2
5.7
Qoa-d
Qoa-e
55,000
± 12,000
RW-11
500,000
±200,000
RW-14
•From the Munsell Soil Color Chart.
TABLE 1. (Continued)
Iron oxide contents and composition (%)
Organic
carbon (%)
Fe 2 0 3 d
Fe 2 0 3 o
Fe 2 0 3 p
FeOT
Ke2Oj
Fe 2 0 3 T
Location and comments
Mouth of Pitman Canyon at Cajon Creek
(granitic debris). 1938 flood deposit
burying the pre-1938 road. Lat.
34°14'32"N; Long. 117°26'23"W.
0.2
0.1
Lone Pine Canyon (Pelona Schist debris).
Two l 4 C dates provide minimum age
constraint; unit predates only last earthquake, providing maximum age constraint. Lat. 34°16'15"N; Long.
117°27'56*W.
2.86
2.44
2.20
2.03
2.09
2.12
2.26
1.77
1.93
1.62
5.22
4.90
4.15
4.13
3.88
0.003
0.01
0.01
0.01
0.01
1.69
0.96
0.99
0.91
0.79
0.52
3.33
3.74
4.10
3.59
2.40
4.40
4.84
5.11
4.47
Mouth of Rat Creek at Cajon Creek
(granitic debris). Age constraints same
as for RW=18. Ut. 34°17'25"N;
Long. 117°26'56"W.
0.15
0.15
0.16
0.16
0.18
0.15
0.01
0.01
0.01
0.01
0.01
0.01
3.17
3.41
2.41
2.69
2.73
2.88
0.78
1.35
1.95
1.65
1.71
1.88
4.30
5.13
4.62
4.64
4.74
5.08
Lone Pine Canyon (Pelona Schist debris).
Incision below Qt-5 isolated Lost Swamp
area from surface flow. Three 14C values
from Lost Swamp sediments date this
event. Ut. 34°16'17"N;
Long. 117°27'56"W.
0.90
0.79
0.17
0.21
0.02
0.02
4.20
3.97
0.86
0.96
5.52
5.37
0.7
0.67
0.21
0.02
2.95
1.65
4.92
0.5
0.63
0.19
0.01
2.61
1.76
4.66
Lone Pine Canyon (Pelona Schist debris).
Offset by the San Andreas fault; age
inferred from offset and slip rate of
24.5 ± 3.5 mm/yr. Absolutely bracketed
by Qt-3 and Qt-5. Ut. 34016'37"N;
Long. 117°28'22"W.
0.62
0.15
0.01
2.72
1.67
4.69
0.51
0.12
0.01
2.11
2.28
4.62
0.69
0.70
0.86
0.81
0.72
0.18
0.17
0.27
0.21
0.21
0.01
0.01
0.01
0.01
0.01
3.00
2.74
2.79
2.86
2.54
1.92
2.36
1.98
2.36
2.26
5.25
5.40
5.08
5.54
5.08
0.72
0.22
0.01
2.46
2.20
4.93
0.63
0.56
0.17
0.13
0.004
0.003
2.41
2.42
2.16
1.85
4.84
4.54
1.8
0.9
0.5
0.81
0.79
0.87
0.85
0.45
0.44
0.23
0.26
0.26
0.26
0.20
0.18
0.01
0.01
0.01
0.01
0.01
0.004
3.94
3.04
2.85
2.88
2.59
2.77
1.07
2.06
2.83
2.15
2.00
1.52
5.45
5.44
6.00
5.35
4.88
4.60
1.6
0.6
0.4
0.2
0.2
1.00
0.89
0.95
1.08
1.26
0.27
0.31
0.34
0.41
0.47
2.07
1.78
1.76
1.72
1.55
2.43
2.27
2.80
2.82
2.67
4.67
4.20
4.71
4.68
4.35
0.3
0.4
0.2
0.2
1.32
1.20
1.09
1.12
0.43
0.43
0.43
039
1.52
1.76
1.60
1.61
3.05
2.61
2.41
2.73
4.70
4.52
4.14
4.48
1.5
0.4
1.50
2.19
2.18
1.50
0.75
0.55
0.37
0.36
0.53
0.58
0.46
0.22
0.18
0.16
2.00
1.03
1.13
1.52
1.74
1.65
1.63
0.5
0.4
1.42
0.63
1.27
0.48
0.88
0.70
0.64
0.49
0.32
0.25
0.10
0.12
0.07
0.12
0.12
0.13
0.16
0.10
0.344
0.133
0.140
0.167
2.9
0.7
0.4
0.2
0.2
0.85
0.69
0.60
0.47
0.46
0.22
0.20
0.24
0.23
0.29
0.7
1.1
0.4
0.2
0.29
0.60
0.58
0.47
0.46
0.11
0.20
0.15
0.14
0.13
0.8
0.5
0.2
0.3
0.55
0.58
0.55
0.59
0.59
0.51
4.3
1.6
2.1
1.2
1.0
0.7
0.5
6.25
7.10
Lone Pine Canyon (Pelona Schist debris).
Lost Swamp sediments were deposited on
Qt-3 as soon as Lone Pine Creek
abandoned the surface; five l4 C dates in
basal clays of Lost Swamp are used to
infer the age of the surface. Ut.
34°16'43'N; Long. 117°28'25"W.
Lone Pine Canyon (Pelona Schist debris).
Minor cut into Qoa-c that appears to be
offset by the San Andreas fault almost as
much as Qt-1. Absolute age limits are
based on the ages of the higher and lower
Qt-1 and Qt-3. Ut. 34°16'6"N;
Long. U7°27'58"W.
Lone Pine Canyon (Pelona Schist debris).
Six C dates in the Qoa-c deposit permit
estimate of surface age based on rate of
fill. C dates in younger units are consistent with extrapolated age; age is
consistent with offset on San Andreas
fault U L 34°16'48"N;
Long. 117028'19"W.
Freeway cut at San Andreas fault (granitic
and gneissic debris). Age based on 1.3- to
1.4-km offset by San Andreas fault, using
slip rate of 24.5 ± 3.5 mm/yr determined
from younger deposits. Similar 0.73-Ma
slip rate justifies extrapolation of the rate
to older deposits. Lat. 34°15'42"N;
Long. 117°26'51*W.
Summit Pass (mixed Pelona + granitic
debris). Unit is incised into Qoa-N (Fig.
2) that contains the Brunhes-Matuyama
polarity reversal (time scale of Harlind
and others, 1982). Age and offset of
Qoa-N yield consistent slip rates on
several faults. Lat. 34°19'I8"N;
Long. I17°25'54"W.
284
14
C dates obtained from the deposits into which
the terraces were cut and from overlying paludal
and colluvial sediments. The amount of offset of
the terraces across the San Andreas fault also
can be used to estimate when particular surfaces
were abandoned by the creek. Establishing the
ages of surfaces is crucial in determining rates of
soil development, Available age control is usually obtained from deposits underlying the surface, and the time when stable surfaces became
established on deposits and the soils began to
form can be only roughly estimated.
The modern Cajon Creek is the result of capture of an older drainage system in the central
Transverse Ranges followed by rapid incision
into Cenozoic sediments. Streams of the older
system flowed north, toward the western Mojave Desert. The beginning of capture, determined on the basis of fossils, paleomagnetic
data, and the offset of key units by faults with
relatively well constrained slip rates, was just
after 0.73 Ma (Weldon and others, 1981; Weldon, 1986). Approximately 500 m of incision
has subsequently occurred in the central part of
the drainage (Fig. 2).
The oldest deposit discussed herein, Qoa-e, is
of middle Pleistccene age (0.5 Ma) and was
formed as the result of a major period of aggradation during the early stages of incision of the
creek. The deposit is dated by its position, early
in the downcutting that began just after a 0.73Ma magnetic polarity reversal (time scale of
Harland and others, 1982) and by its offset of
~ 1 km by the Cleghorn fault, which has a slip
rate of 2 mm/yr (Weldon and others, 1981).
The 0.73-Ma magnetic polarity reversal occurs
in a deposit (Qoa-N) that has been correlated
with the older alluvium of Noble (1954) and
that is the youngest deposit predating downcutting of Cajon Creek. The next youngest major
datable deposit, Qoa-d, is of late Pleistocene age
(-55,000 yr old). It is dated by its 1.3-km offset
across the San Andreas fault, which has a slip
rate of 25 mm/yr (Weldon and Sieh, 1985).
Formation of Qoa-d was the result of as much
as 85 m of aggradation. Reincision into this
massive deposit created the "inner gorge" of
Cajon Creek (Fij;. 2). Within the inner gorge,
there are a latest Pleistocene depositional terrace
(Qoa-c) and a late Holocene depositional terrace (Qoa-a). A total of seven terraces have been
recognized in the inner gorge (Fig. 2, Qt-l-Qt7), two on the surfaces of the depositional terraces Qoa-c and Qoa-a and five erosional terraces formed during latest Pleistocene and
Holocene time. This study focuses on the soils
formed on the dated terraces of the inner gorge;
McFADDEN AND WELDON
however, soils developed on the surfaces of
Qoa-d and Qoa-e are included for comparison
with the younger soils.
The depositional-terrace deposits in the inner
gorge have yielded 18 radiocarbon ages from
both within and, locally, at the top of them
(Weldon, 1986). Rates of sedimentation estimated from radiocarbon ages in the fill can be
extrapolated to estimate the age of the top of the
fill. Minimum ages of the terrace surfaces can
also be determined on the basis of the age of
sediments subsequently deposited on the cut terraces. For example, Qt-3 is dated on the basis of
14C
ages obtained from sediments of Lost
Swamp (Fig. 2), as described in more detail in
Weldon and Sieh (1985) and Weldon (1986).
After a terrace surface was abandoned by Cajon
Creek, offset across the San Andreas fault began
to accumulate. Given the slip rate of the fault
(Weldon and Sieh, 1985), the offset of a terrace
by the fault provides an accurate way of estimating when the terrace was actually abandoned
and when soil formation was initiated. After
abandonment of a terrace, rapid incision tended
to isolate the broad, low-gradient terrace and
prevented significant degradation of, or deposition on, the terrace.
Locally, soil-stratigraphic data show that
some of the terraces had been subjected to previously unrecognized erosion or colluviation,
chiefly at sites near hillslopes. Only soils from
the geomorphically most stable sites were selected for detailed textural and chemical analysis
(Table 1). Morphological data from these soils
and less detailed data for the other soils are included in Tables A and B,1 which are on file
with the Geological Society of America Data
Repository.
FIELD AND LABORATORY
METHODS
sodium pyrophosphate, wet-sieve separation of
the sand and silt + clay fractions, and pipette
extraction for clay content. Organic material in
A, AC, and the upper part of the B horizons was
removed prior to particle size analysis by H2O
digestion. Organic-carbon content in these horizons was measured colorimetrically using a
Bausch and Lombe Spectronic 20 spectrophotometer,2 after the method described by Metson
and others (1979). Soil pH was measured in
1:10 soiil-to-water ratio in 0.01 M CaCl2.
In well-drained, oxidizing soil environments,
ferrous iron is progressively converted to essentially insoluble ferric-iron oxides that accumulate in increasingly older and typically redder
soils (Schwertmann and Taylor, 1977). Changes
in soil iron oxide content and composition are
closely related to degree and nature of soil development. Several methods, thus, were used to
evaluate soil iron. Total soil iron (represented as
Fe203T) was extracted by using the hydrofluoric, nitric, and perchloric acid digestion method
(Husler, 1969). Hydrofluoric and sulfuric acid
digestion, followed by potassium dichromate titration, was used to determine ferrous iron
(FeOT) (Kolthoff and Sandell, 1961). The difference between FeOT and Fe203T determines
the ferric component (Fe203) of total soil iron.
Comparisons of Fe20 3 , FeOT, and Fe2C>3T
data among soil horizons allow us to estimate
the degree of chemical alteration and relative
losses or gains in iron content due to soil formation. Extraction of total ferric iron present in
oxyhydroxide phases such as hematite and goethite (Fe20 3 d) was accomplished by using the
dithionite-citrate-bicarbonate procedure of
Mehra and Jackson (1960). Ferric iron present
in poorly crystalline oxyhydroxide phases
(chiefly ferrihydrite) and organic complexes
(Fe20 3 o) was extracted by using the oxalate
extraction method of McKeague and Day
Soil profiles from hand-dug pits were described wherever possible. Because the depth of
weathering that is associated with many soils
exceeds 2 to 3 m, the deepest soil horizons described were in stream cuts or in road cuts, and
the upper horizons were matched with horizons
in pits. Soil profiles were described and sampled
primarily according to the procedure and terminology of the Soil Survey Staff (1951, 1975).
Particle size distribution was determined by
clay dispersal of the <2-mm fraction in 10%
(1966). Because magnetite, which varies in these
soils between 0.19 and 0.38 wt%, is slightly soluble under the conditions of oxalate extraction
(Rhoton and others, 1981; Walker, 1983), it
was removed prior to Fe 2 0 3 o extraction by
using a strong magnet. Iron present in organic
complexes (Fe203p) was extracted from selected soils by using the method of McKeague
(1967). Determination of Fe20 3 d, Fe2030, and
Fe20 3 p provides additional data for evaluating
the magnitude of chemical alteration of ironbearing minerals and of gains in iron oxides due
'Tables A and B may be obtained free of charge by
requesting Supplementary Data 87-07 from the GSA
Documents Secretary.
2
Use of trade names in this paper is for descriptive
purpose:! only and does not constitute endorsement by
the U.S. Geological Survey.
SOIL DEVELOPMENT ON QUATERNARY TERRACES, CALIFORNIA
A
10
E
u
0
,_
10
SILT (%)
10
o
10
ml
I
10
20
I
L_
20
?
2040
T
0Lil
Q
60
40 yrs. B.P.
275 yrs.
B.P
5900
285
yrs. B.P.
7100 yrs.
8300
B.P.
yrs. B.P.
Figure 3. Estimated increase in pedogenic silt content (A silt %) in soils formed on
terraces of Cajon Creek, showing initially
rapid accumulation of silt in Holocene soils. A
silt % determination is based on maximum silt
content of the unaltered or least altered C
subhorizon. Minimum increase in silt % considered pedogenic in origin = +3%.
A SILT (%)
0
10
,
20
,
3,0
o
20
I
X
(a.
HI
Q
40
60-
80
11,500 yrs.
1 2 , 4 0 0 yrs.
B.P.
to additions by other processes. Extracted iron
(Fe2C>3T, Fe2C>3d, Fe20 3 o, Fe 2 0 3 p) was measured by using a Perkin and Elmer 303 atomic
absorption spectrometer.
In order to determine the absolute amount of
increase of an iron component in a given soil,
the amount of the iron component in the unaltered parent materials must be determined. On
middle and late Holocene terraces, unaltered
parent materials are present, although mottling
of the matrix and partly grussified stones in
some cases are present and demonstrate that
minor chemical alteration has taken place locally. Unaltered parent materials of early Holocene terraces are not present in the upper several
metres, and completely fresh parent materials of
latest Pleistocene terraces are not present in the
upper several metres. Completely fresh parent
materials of late to middle Pleistocene terraces
probably no longer exist. As there are few sedimentological differences among different fluvial
deposits, parent material characteristics for early
Holocene and Pleistocene soils in the study area
can be estimated on the basis of parent material
data for younger Holocene soils. The gain in an
iron component for the entire soil (profile content) is calculated by summing the net increases
of the component (weight percent of iron component in a measured horizon minus that in the
parent material x horizon thickness) of all horizons above the shallowest unoxidized or least
oxidized horizon.
B.P.
SOIL DEVELOPMENT ON
HOLOCENE TERRACES
Morphology, Chemistry, and Mineralogy
The initially developed and most prominent
soil horizon that occurs on late Holocene geomorphic surfaces is the darkened A horizon
(Table 1). This horizon reflects the rapid establishment of the dense chaparral vegetation
community on abandoned flood plains. In late
Holocene soils on Qt-6, a thin (7 to 32 cm),
slightly reddened horizon (Cox) is always present below the transitional AC horizon. Reddening is due partly to ferric oxide stains but is also
due partly to reddish particles of silt that coat
larger skeletal grain surfaces. The presence of
such silt coatings (siltans) indicates downward
translocation of silt. Stones in the late Holocene
soils are not visibly weathered, as indicated by
the sharp ring from a hammer blow, the absence
of weathering rinds, and the presence of smooth,
stream-worn surfaces.
Soils on middle and early Holocene surfaces
(Qt-5, Qt-4, Qt-3) possess even thicker, darker
epipedons (mollic A horizons) and reddened,
silt-enriched, color B horizons (one type of Bw
horizon) (Tables 1 and A3). The thickness of the
A horizon generally ranges from 26 to 35 cm,
although a very thick mollic horizon associated
3
See footnote 1.
with a soil profile that formed on an -2,000-yrold fan deposit near the study area is 80 cm
thick (profile RW-16, Table A). The uppermost
soil horizons contain large amounts of silt and
organic carbon (Table 1, Fig. 3). The silt occurs
mainly in the soil matrix, but some silt and organic matter are also present as thin coatings on
stones. The color B horizon is present below the
transitional AB or BA horizon. Significant silt
has accumulated in the B horizon (Table 1, Fig.
3), some of which is present as coatings of silt on
skeletal grains. If conventional techniques of
particle size analysis are used, very little clay is
detectable in the B horizon; however, micromorphologic evidence shows that a very small
amount of clay is probably present, occurring as
colloidal stains on skeletal grains. In marked
contrast to stones in late Holocene soils, some of
the schistose and many of the coarse-grained
plutonic stones in middle and early Holocene
soils have been weathered to a grus or near-grus
state.
Holocene soils acquire slightly to moderately
acidic pH values. The pH is lowest in the A
horizon and increases significantly with depth
(Table 1).
The maximum content of organic carbon occurs in the A horizon. In the middle Holocene
soils, significant amounts of organic carbon are
also present in transitional horizons of the upper
part of the Bw or Bt horizon. The maximum
total organic carbon (profile carbon) occurs in
middle Holocene soils (Fig. 4).
Significant quantities of iron oxides have accumulated in Holocene soils (Table 1). Gains in
Fe203d and Fe2C>3T content are most pronounced initially in the A horizon, but increasing amounts of these constituents occur in the
Bw horizons of early Holocene soils (Table 1).
In the soil on the Qt-5 surface (RW-12), the
FeOT and Fe2C>30 content and the Fe2C>3d and
Fe203T content of the Cu/Cox horizon are uniformly higher than those of the unaltered parent
286
McFADDEN AND WELDON
— Organic Carbon
— Epipedon Thickness
Time ( y e a r s
Figure 4. Changes in profile organic carbon content and epipedon thickness as a function
of soil age in the Cajon Pass area.
materials of other Holocene terraces or in the
least altered materials at the base of Qoa-c terrace fill. This increase indicates that iron enrichment is associated with stratigraphic variation or possibly with ground-water alteration
and is not attributable to pedogenesis. The increase in Fe2C>3d and Fe2C>3T due to soil development thus is probably much greater than is
apparent. The trends of ferric-iron (Fe2C>3) accumulation in Holocene soils are similar to the
trends observed for Fe203d and Fe203T, except
for a usually low Fe2C>3 content in the A horizon (Table 1). The data also show that part of
the accumulated Fe2C>3 is due to causes other
than simple increases in ferric iron present as
ferric oxyhydroxides. For example, the increase
in Fe2C>3 percentage in the 2Bwl horizon of the
soil on the Qt-3 surface (RW-13) relative to the
Fe2C>3 percentage in the soil parent material is
~ 1.0%, compared i:o an Fe2C>3d increase of only
0.3% to 0.4%. Gains in Fe 2 0 3 thus might be
partly due to addition of minerals that initially
have ferric iron, such as biotite or some clay
minerals. Changes in FeOT content are generally similar to changes in Fe 2 03d and Fe203T
content in Holocene soils, although the magnitude of the increase of FeOT below the A horizon relative to FeOT in soil parent materials is
usually much smaller than that for associated
increases in Fe20;,d, Fe203, and Fe203T content (Table 1). Increases in Fe2O30 content are
small compared to those observed in other iron
components (Table 1). As in the case of Fe203,
such gains typically occur in the B horizon
rather than in the A horizon. Low Fe203P con-
tents indicate that very little of the ferric iron in
Holocene soils is organically complexed and
that content of organically complexed iron does
not increase with soil age (Table 1). Increases in
Fe2030 content are thus attributable largely to
the accumulation of ferrihydrite. The presence
of at least some organically bound iron, in association with the gains in organic-carbon content,
however, is consistent with the edaphic environment (extensive adsorption of organic complexes) proposed by Schwertmann and Taylor
(1977) that favors formation and prolonged stability of ferrihydrite.
Influence of Eolian Dust Influx
Two of the most important aspects in soil
development during the late and middle Holocene are the initially rapid accumulations of silt
and of organic matter (Figs. 3 and 4, Table 1).
Field observations, micromorphic studies, and
laboratory analyses indicate that almost no
chemical and physical weathering of the parent
materials of soils has occurred on late and middle Holocene terraces. Field observations also
show that grussification of boulders occurs in
place and produces little matrix that is finer than
sand. These observations indicate that much of
the silt accumulating in Holocene soils is contributed as eolian dust. Eolian processes are consistent with the abundant coatings of silt present
on skeletal grains and stones, an association that
indicates downward translocation of silt from
the surface horizons.
Many previous studies have demonstrated
that eolian processes can contribute silt, clay,
and calcium carbonate to soils. Other studies
have demonstrated that the source of dust can be
either local (Lattman, 1973; Peterson, 1980;
McFadden and others, 1986) or thousands of
kilometres distant (Prospero and others, 1981;
P6we and others, 1981; Muhs, 1983). The Mojave Desert, located seasonally upwind of the
study area, probably serves as a primary source
of eolian dust in the Cajon Creek area. Evidence
for eolian activity in the area includes the presence of ridge dunes along the northern edge of
Cajon Pass and the extensive eolian deposits in
the San Bernardino Valley to the south.
Changes in the iron oxide content and composition of soils on Holocene terraces as depth
and soil age increase also indicate that many of
the accumulated iron oxides have been derived
primarily by incorporation of eolian materials
rather than by chemical alteration. For example,
despite the well-drained, oxidizing, and acidic
environment of Holocene soils, ferrous-iron content and the Fe0T/Fe203 ratio decrease with
depth (Fig. 5). Furthermore, the very minor increases in ferrihydrite relative to significant accumulations of more crystalline ferric-iron
oxides in Holocene soils also suggest that little
chemical alteration of ferrous-iron to secondary
ferric-iron oxides has occurred. This pattern of
iron accumulation in Holocene soils is therefore
attributed to incorporation of iron-bearing eolian dusl: at a rate that exceeds the rate of chemical alteration of ferrous-iron to secondary
ferric-iron oxides. Ferrous iron, present in ferromagnesian minerals, magnetite, and various other iron oxides, is a moderately abundant constituent of eolian dust (Pewe and others, 1981).
The relative depletion of ferrihydrite and the
low Fe20 3 o/Fe203d ratio of the uppermost A
horizons of Holocene soils also suggest that the
eolian dust transported through Cajon Pass has
less ferrihydrite than do the soil parent materials.
Statistical analysis of iron content and silt
content of the soil supports the hypothesis of
eolian iron addition. Use of the nonparametric
Spearman rank correlation test shows that the
secondary iron oxide content and the silt content
of Holocene soil horizons in the study area (n =
30) are positively correlated; correlation is significant at the 99.5% level of confidence (a
= 0.05). Ordinary least-squares regression analysis strongly indicates a linear relation between
these two components.
Fe 2 0 3 d % = 0.01 silt % + 0.43, r = 0.83
Similarly, analysis of ferrous-iron and silt content, using the Spearman test, shows that these
SOIL DEVELOPMENT ON QUATERNARY TERRACES, CALIFORNIA
FeOT/ Fe203
1.0
=
25
2.0
3.0
5.0
J
.J
I
E
LEGEND
ftl
Q_
0 >
O
4.0
-
2 7 5 yr. B. P
2 7 5 yr. B.F?
5 , 9 0 0 yr. B P
7 5
7,100 yr B P
8 , 3 0 0 yr. B. P
100
11,500 yr. B P
1 2 , 4 0 0 yr. B P
125
Figure 5. Changes with depth in the ratio of total ferrous iron to total ferric iron (FeOT/
Fe203) for Holocene and latest Pleistocene soils in the Cajon Pass area. The ratio increases in
the A, AC, or B horizon relative to the unaltered or least altered C horizon in all soils.
two components are positively correlated (a =
0.025) in most Holocene soils. Simple linear
regression analysis of these two components also
indicates correlation of the two components.
A FeO % = 0.04 A silt % + 0.04, r = 0.57
Logarithmic transformation of silt content improves the degree of correlation.
A FeO % = 0.16 In A silt % -3.66, r = 0.79
Statistical analysis indicates functional dependence of the ferrous-iron and ferric-iron oxide
content on the silt content, a relation consistent
with addition of iron-bearing silt to Holocene
soils.
Slight reddening of subhorizons in Holocene
soils demonstrates that at least some authigenic
ferric-iron oxides have accumulated because
reddened soils require formation of minerals
such as hematite and ferrihydrite (Schwertmann
and Taylor, 1977; Childs and others, 1979;
Schwertmann and others, 1982). The relatively
low Fe203 content of the uppermost A horizon
is consistent with chemical alteration and
downward translocation of ferric iron. As these
and others studies (Walker, 1967) have shown,
however, very little authigenic iron oxide accumulation is required to make soils or sediments considerably redder than the initial color
of the unaltered materials. For example, increases of 0.02% to 0.07% ferrihydrite appear to
be sufficient to change the initial olive-gray color
of the parent material to a light brownish-gray
or yellowish-brown color (Table 1).
SOIL DEVELOPMENT ON
PLEISTOCENE TERRACES
Morphology, Chemistry, and Mineralogy
Soils on the Pleistocene terraces are morphologically much better developed than are Holocene soils, possessing argillic B horizons that
become progressively thicker, redder, and more
clay rich with increasing age (Table 1). A significant amount of the clay in the argillic horizon is
illuvial, occurring as coatings which thicken increasingly with time on the grains or peds, as
pore-filling material, or as bridges. In late and
middle Pleistocene soils (RW-11, RW-14),
stones are rare or have been completely weathered to a grus-like state. The epipedon thickness and organic matter content of the late
Pleistocene soil are less than the epipedon thickness and profile organic matter content of most
Holocene soils (Fig. 4). Exposures of the deposits underlying Pleistocene surfaces show that the
Cox horizon extends to depths of at least 5 m.
Alteration to depths of 20 to 30 m is indicated
by slight reddening of the matrix or the presence
of locally reddish mottles.
The oldest surface (Qoa-e) is generally heavily dissected. We described the soil on the most
stable remnant of the surface at a site where the
soil is buried by ~ 1 m of eolian sand that has a
weakly developed soil. Although we cannot be
absolutely confident that no stripping of the argillic horizon has occurred, the argillic B horizon
is characterized by the maximal reddish color,
clay content, and structure observed in the study
area. In part of the argillic horizon (2Bt3),
287
mottled and bleached appearance suggests that
for unknown reasons, the soil was subjected to
reducing conditions after formation of much of
the argillic horizon had occurred. Chemical and
mineralogical evidence do indicate some net loss
of Fe 2 03d from this horizon (Table 1). The argillic B horizon is at least 3.35 m thick, and
exposures in road cuts suggest that the horizon
may actually be considerably thicker. Weakly
altered parent materials that exhibit colors
corresponding to Cox horizon colors of younger
soils were observed only at a depth of 15 m.
The maximum Fe203d, Fe2030, and Fe203
contents of latest Pleistocene soils, in contrast to
most Holocene soils, occur in the argillic B horizon and not in the A horizon (Table 1). Profile
Fe203d and Fe 2 0 3 o contents are also much
greater than in Holocene soils. Maximum FeOT
content, however, as in the case of Holocene
soils, occurs in the A horizon and decreases rapidly with depth, although the minimum FeOT
content of the soil on the Qt-1 surface occurs in
the B horizon and not in the C horizon. In most
respects, changes in Fe203p content and soil pH
of latest Pleistocene soils resemble those recognized in middle and early Holocene soils.
The maximum Fe 2 03d and Fe203 contents
of late and middle Pleistocene soils, as in the
case of latest Pleistocene soils, occur in the argillic horizon (Table 1). In the late Pleistocene soil
on Qoa-d, maximum Fe 2 0 3 o content also occurs in the argillic horizon, the minimum FeOT
occurs in the argillic horizon, and Fe203p content is uniformly low (Table 1). In the middle
Pleistocene soil on Qoa-e, the minimum FeOT
content also occurs in the argillic horizon, although in contrast to younger Pleistocene soils,
the Fe 2 030 content is also very low in the argillic horizon.
Formation of Authigenic Ferric-Iron Oxides
Several aspects of soil development on increasingly older Pleistocene terraces contrast
significantly with those of Holocene soil development: an argillic B horizon appears, the A
horizon declines, and the ferric oxides and clay
accumulate, whereas the ferrous iron in the argillic horizon becomes depleted. The pattern of
iron oxide formation in Pleistocene soils strongly
indicates that the increasingly larger quantities of
ferric-iron oxides in the argillic horizon are due
to chemical alteration and formation of authigenic ferric-iron oxide minerals rather than to
incorporation of iron-bearing minerals present
in eolian dust.
Evidence of in situ chemical alteration is provided by systematic changes in the Fe20 3 d/
288
McFADDEN AND WELDON
Fe203d /
0.1
0.2
hp
25
Fe203T
0 3
0.4
0.5
Íj
II
J!
50
75
Q.
a> 1 0 0
O
¡
LEGEND
2 7 5 yr. B P
125
5 , 9 0 0 yr. B P
8,350 yr.BP
I I , 5 0 0 yr.B.P
150
1 2 , 4 0 0 yr.B.P
5 5 , 0 0 0 yr. B P
5 0 0 , 0 0 0 yr. B.P
175
Figure 6. Changes with depth in ratio of
total iron oxyhydroxides to total iron
(Fe20 3 d/Fe20 3 T) in soils of the Cajon Pass
area. Significant increase in this ratio occurs
in the argillic horizon of latest Pleistocene
and older soils.
Fe 2 0 3 T, Fe0T/Fe 2 0 3 , and Fe 2 0 3 o/Fe 2 0 3 d
ratios. Increases in the Fe20 3 d/Fe20 3 T ratio
(especially in B horizons) are closely related to
an increasing degree of primary-mineral alteration, soil development, and soil age (Rebertus
and Buol, 1985), In Holocene soils of Cajon
Creek, only slighl. increases in the ratio are observed (Fig. 6). In increasingly older Pleistocene
soils, however, this ratio increases significantly
in the B horizon and attains maximum values in
the middle Pleistocene soil. The progressive increase in the Fe20 3 d/Fe20 3 T ratio shows that
an increasingly larger proportion of ferrous iron
in primary minerals of parent materials or in
incorporated eolian dust has been converted to
authigenic ferric-iron oxides. This conversion is
consistent with i:he progressively marked decreases in the Fe0T/Fe20 3 ratio in subhorizons
of the argillic horizon compared to that ratio in
the parent material of Pleistocene soils (Figs. 5
and 7). Increases in the F e 0 T / F e 2 0 3 ratio in
the A or uppermost B horizons of Pleistocene
soils indicate continuing incorporation of ferrous
iron in eolian dust; however, the rate of chemical alteration of ferrous to ferric iron in the maximal argillic horizon exceeds the rate at which
dust-derived ferrous iron is added.
Changes in the Fe 2 0 3 o/Fe20 3 d ratio also
provide evidence for formation of authigenic
ferric-iron oxides in soils on Pleistocene terraces.
Increases in profile Fe20 3 o content in soils on
increasingly older Pleistocene terraces (Fig. 8)
indicate that part of the depleted ferrous iron can
be attributed to alteration of ferrous-iron-bearing minerals to ferrihydrite, causing an increase
in the profile Fe 2 0 3 o/Fe20 3 d ratio. Ferrihydrite, however, is a metastable mineral and
eventually transforms to more stable minerals,
primarily hematite (Schwertmann and Taylor,
1977). The value of the profile Fe20 3 o/Fe20 3 d
ratio thus cannot possibly continue to increase
with soil age; after some duration of soil development, the ratio should achieve a maximum
value and decline thereafter in soil-forming environments that ultimately cause transformation
of ferrihydrite to more crystalline hematite
(McFadden and Hendricks, 1985). In the Cajon
Creek chronosequence, the ratio is greatest during the latest Pleistocene and decreases after that
despite continuing formation of ferrihydrite
(Fig. 8).
The net depletion of Fe20 3 o in the argillic
horizon of the middle Pleistocene soil (Table 1,
Fig. 8) results in a negative profile ratio of
Fe20 3 o/Fe20 3 d and indicates that nearly all
accumulated ferrihydrite has been converted to
a more crystalline iron oxide. The increasingly
bright color of these soils indicates that this mineral is very likely hematite. This conversion
probably accounts for the strong relation between soil color and Fe20 3 d content of soils in
the Cajon Creek area (Table 1). The increasing
abundance of pedogenic hematite with increasing soil age has been documented in many studies (Childs and others, 1979; Torrent and others,
1980; Schwertmann and others, 1982) and apparently accounts for the increasingly bright red
colors of such soils. Using Hurst's (1977) redness
index, R = (Hue * Value)/chroma, McFadden
and Hendricks (1985) demonstrated that the
linear relation between R and Fe20 3 d percentage is statistically significant for early Holocene
and Pleistocene soils throughout the Transverse
Ranges. Statistical analysis of R (modified to
account for 5Y and 2.5Y hues), as well as of
another index of reddening and brightening
(rubification) defined by Harden (1982), and
Fe20 3 d content using both nonparametric
(Spearman rank correlation) and ordinary least
squares regression tests shows that soil reddening
correlates very strongly with Fe20 3 d in early
Holocene and older soils (minimum a = 0.025;
range in value of r = 0.76 to 0.92). The strong
correlation of soil color and Fe20 3 d content
Fe0T/Fe203
0
0.1
0.2
0.3
0.4
0.5
0.6
0.7
IT
r~
LEGEND
25
—
50
5 5 , 0 0 0 yrs. B P
5 0 0 , 0 0 0 yrs. B P
.Composition of
I least a l t e r e d
75
horizon
100
I
§"125
150
I
"1
I500
1525
Figure 7. Changes with depth in the ratio
of total ferrous iron to total ferric iron
(Fe0T/Fe 2 0 3 ) for late and middle Pleistocene soils in the Cajon Pass area. Major decrease in the ratio occurs in the maximal part
of the argillic horizon. Arrows show this ratio
calculated on basis of data for least altered
horizon of each soil; deposits of the terraces
on which the soils occur have been altered to
a depth exceeding 20 m.
supports the chemical evidence for the formation of a greater relative abundance or proportion of ferric-iron oxides in Pleistocene soils
discussed above.
The growth of the argillic horizon that is increasingly rich in authigenic iron oxides is accompanied by thinning of the A horizon. This
process cannot be attributed to erosion or stripping of a once thicker A horizon. Latest Pleistocene terraces that are completely isolated from
fluvial flow exhibit little evidence of erosion.
The remnants of the original bar-and-swale topography exposed in cliffs and in our pits and
the lack of a mechanism to move large boulders,
in some cases exceeding 2 m in intermediate
diameter, off the wide, undissected terraces
clearly indicate that the surface has not been
stripped. Surficial erosion of the finer material
would progressively expose boulders; instead,
these boulders are gradually being buried. Field
observations show that the bar-and-swale topography of a recently abandoned terrace is
gradually eliminated, primarily by infilling of
swales with eolian material and with material
off the bars by sheetwash and bioturbation.
These processes produce smooth surfaces on terraces as young as the early Holocene. Locally,
SOIL DEVELOPMENT O N QUATERNARY TERRACES, CALIFORNIA
289
0.50r
o
ro
"O
ro o v s
<_> O
CM
CM
£ £
0
< <
W - 0.25
Profile ratio
-0.50 •
500
Figure 8. Changes in Harden (1982) soil index, profile ferric-iron
oxyhydroxides (Fe 2 0 3 d), profile poorly crystalline ferric-iron oxyhydroxides (Fe 2 0 3 o), and profile Fe2C>30/Fe2O3d ratio as a function of
soil age in the Cajon Pass area.
400
400 -
E
0
1
300
LEGEND
300
—
3
Q.
E e 2 0 3 o , Profile
CD
X
ro
O
<u°
• F e 2 0 3 d , Profile
200
200
• Profile Index
LL
a)
o
CL
100
100
-
100
1,000
10,000
Time ( years)
100,000
TABLE 2. SOIL PROFILE INDICES AND THEIR WEIGHTED
MEANS FOR HOLOCENE AND PLEISTOCENE SOILS
IN THE CAJON PASS AREA, SOUTHERN CALIFORNIA
the relatively thick, silt-rich and nongravelly A
horizons of early Holocene and latest Pleistocene soils are probably the result of cumulative
soil development in former swales.
The declining thickness of the A horizon must
be attributed to upward thickening of the B horizon and to a concomitant increase in the magnitude of oxidation of organic matter at the
expense of the A horizon (McFadden and Hendricks, 1985). Increases in water-holding capacity and decreases in infiltration rates favor an
increase in the magnitude of oxidation of soil
organic matter, particularly during hot, dry
summer months. Decreases in permeability of
the soil also may inhibit mechanical translocation of large fragments of undecayed to partly
decayed organic matter that readily accumulate
in loose, gravelly Holocene soils.
RATES OF SOIL DEVELOPMENT
Soil Morphology and Iron Oxide Content
Two soil characteristics that have systematically increased during the past 0.5 m.y. in the
Cajon Creek area are soil morphologic development and ferric-iron oxide content (Figs. 6,7,
and 8; Table 1). The over-all degree of morpho-
logic development of a given soil can be quantified by using an index derived by Harden (1982)
that combines horizon thickness and other field
properties. The value of this index has been
shown to increase systematically with soil age in
many chronosequences in diverse climates and
parent materials (Harden and Taylor, 1983;
McFadden and others, 1986; Ponti, 1985). The
soil profile indices of selected soils in the study
area (Table 2) correlate strongly with soil age (a
= 0.005, Spearman rank correlation). Ordinary
least-squares regression analysis of the indices
and soil age yields the following equations.
(1) Profile index = 0.001 + 12.11 (terrace age), r
= 0.998, and (2) log (weighted mean) = 0.093 0.15 (terrace age), r = 0.957 (weighted mean is
the profile index/profile thickness). Similarly,
profile Fe20 3 d (pF) is also strongly correlated
with soil age (a = 0.005, Spearman rank correlation). The rate of pF increase with age, estimated by using ordinary least-squares regression
analysis, is (1) pF = 121.2 log (terrace age) 378.9, r = 0.80, and (2) log (pF) = 0.62 log
(terrace age) - 0.92, r = 0.94. In order to test
further the use of these soil parameters to estimate terrace age, several soils must be analyzed
on each terrace, thereby permitting determination of the degree of variability of a given pa-
Soil
profile
Age
(yrB.P.)
Profile
index*
Weighted
mean
RW-9
1947
2.97
0.06
RW-10
275
•385
-75
5.49
0.11
RW-18
275
+385
-75
6.93
0.10
RW-12
5900
±900
13.65
0.14
RW-15
7150
±1200
14.60
0.15
RW-13
8350
+900
-500
13.04
0.13
RW-6
11,500
+2000
-3000
19.12
0.19
RW-17
12,400
±1000
20.11
0.17
RW-ll
55,000
±12,000
50.97
0.24
RW-14
500,000
±200,000
277.99
0.40
Note: parent-material properties for soils that are rich in Pelona Schist: dry
color, 5Y 6/2; moist color, 5Y 4/2; texture = gravelly sand; structure = single
grain; dry consistence = loose; wet consistence - nonsticky and nonplastic; clay
films = none; pH = 7.0 Parent-material properties for soils that have little Pelona
Schist are the same except for the dry color, 10YR 6/3, and moist color, 10YR
4/3. Maximum values for parameters used to calculate the profile index for
soils in this study were calculated on the basis of morphologic data for soils
reported by Harden (1982) except for soils formed in parent materials of Pelona
Schist, in which maximum value of rubification equals 220 points.
•Soil profile index (Harden, 1982) determined for 100-cm depth except for
RW-I7 (110 cm), RW-I1 (200 cm), and RW-14 (701 cm), calculated on the
basis of properties indicated in note.
290
McFADDEN AND WELDON
A CLAY (%)
5
10
5
10
0
5
1
10
1
15
I
20
1
25
L
Figure 9. Changes with depth in pedogenic clay content (A clay %) in soils formed
on Holocene and late Pleistocene terraces of
Cajon Creek. A clay % determinations on the
basis olf maximum clay content of least altered C subhorizon.
7 1 0 0 yrs. B.P.
8 3 0 0 yrs. B.P.
1 1 , 5 0 0 yrs. B.P.
rameter for a given terrace. On the basis of data
for the Cajon Creek chronosequence, we conclude that soil morphology and iron oxide content are potentially excellent indicators of
absolute age of Quaternary deposits in much of
the Transverse Ranges over a span of 0.5 m.y.
Clay Content
Pedogenic clay content also generally increases with soil age in the Cajon Pass area, a
feature noted elsewhere in the Transverse
Ranges (Keller and others, 1982) and in many
other areas (for example, Bockheim, 1980; Gile
and others, 1981; Marchand and Allwardt,
1981; Guccione, 1985; McFadden and Bull,
1987). In contrast to these studies, however, clay
content increased very little during the first
8,300 yr of soil development (Fig. 9). A relatively sudden increase in the rate of clay accumulation during the subsequent 3,000 yr of soil
development is required to create the argillic horizon of latest Pleistocene soils.
Such an apparent change in the rates or processes of soil development has often been
attributed to the significant changes in climate
that have occurred during the Quaternary (Hunt
and Sokoloff, 1950; Morrison and Frye, 1965;
Yaalon, 1971; Gile and others, 1981; Chartres,
1980). It is tempting to ascribe the presence of
increasing amount!! of ferric-iron oxides in argillic B horizons of latest Pleistocene soils and the
lack of such horizons in early and middle Holocene soils to changes in climate and, by inference, to changes in rates and processes of soil
development that occurred at the end of the
Pleistocene. Deposition of unit Qoa-c and its
subsequent incision, however, was probably
1 2 . 4 0 0 yrs. B.P.
5 5 . 0 0 0 yrs. B.P.
triggered by this climatic change (Weldon,
1983,1986); thus, soil development on the latest
Pleistocene surfaces has occurred almost entirely
during the Holocene.
Even in the unlikely case that the climate did
not change until the Holocene, soils on the latest
Pleistocene erraces could not have developed
an argillic horizon in coarse porous gTavels during the 1,500 to 2,400 yr of Pleistocene climate
that the soils experienced. The climatic changes
during the Holocene certainly have not been
nearly as significant as the Pleistocene-toHolocene change; therefore, the rapid development of the argillic horizon that is observed in
the study area cannot be attributed to climatic
change from relatively moist conditions, favoring rapid rates of chemical alteration, to drier
conditions, presumably favoring much lower
rates.
We attribute the relatively sudden appearance
of the argillic horizon to the significant increase
in silt content and its impact on soil permeability
and water balance. Although eolian dust contains clay, apparently little eolian clay is entrapped in gravelly sediments that have very
high permeability. Most of the clay is translocated out of the zone of A and B horizon development, as shown by clay-bearing coatings on
stones at depths that exceed 5 m in Holocene
deposits and by the very low clay contents in
young soils. High soil permeability also decreases the time during which soil moisture is
retained and thereby limits the magnitude of
chemical weathering. While silt and organic
matter continued to accumulate in the upper
profile, however, conditions favoring clay entrapment and chemical alteration develop as the
initially noncolloidal soil pores are gradually
filled with silt and organic matter. For example,
an increase in the silt content from 1% to 18%,
with no change in clay content, theoretically increases available water-holding capacity (AWC)
(Birkeland, 1984) by as much as 35% and concomitantly lowers infiltration rates. Increasing
the silt content to 28% further increases AWC
by another 37%. The concomitant accumulation
of organic matter probably also increases AWC;
thus, AWC is potentially doubled by the accumulations of silt and organic matter measured
in the Cajon Pass area.
Latest Pleistocene to early Holocene eolian
influx rates may have been greater than subsequent influx rates, or eolian dust during the latest
Pleistocene to early Holocene might have contained more clay than did the subsequent eolian
dust. Either of these factors would accelerate the
rate of argillic horizon development, although
an initial period of accumulation of material that
reduces soil permeability still would probably be
required. Therefore, the formation of the argillic
horizon on the latest Pleistocene terraces certainly took place during the late Holocene, after
a system that could hold the clay and produce
higher AWC had evolved. Older Pleistocene
soils presumably have also passed through this
threshold.
It is difficult to distinguish clay produced by
weathering from clay produced by addition of
eolian material in the Pleistocene argillic horizons. Alteration of ferrous iron in fine-grained
minerals, such as biotite in the parent materials,
or in eolian dust certainly may result in in situ
formation of authigenic clay minerals. McFadden and Hendricks (1982), for example, reported that systematic increases in vermiculite
content occur in early Holocene to late Pleistocene soils on fluvial deposits throughout the
Transverse Ranges. Although vermiculite may
be present in dust, the increasing abundance of
this mineral relative to other clay minerals is
more likely caused by alteration of appropriate
mafic minerals to vermiculite and iron oxides.
SOIL DEVELOPMENT ON QUATERNARY TERRACES, CALIFORNIA
TABLE 3. PROFILE DATA FOR CHRONOSEQUENCE OF SOILS IN THE MERCED AREA, CALIFORNIA
Formation
or
unit
Surface age
(yr)
Profile
Generalized
horizon
sequence
Parent
materia]
Profile
depth
(cm)
Profile
index
Weighted
mean
Post-Modesto
3,000
PM8
A/C
fSL/SiL
76
13.10
0.17
Post-Modesto
3,000
PM16
A/Cox/C
fSL/SL
236
16.28
0.07
Modesto,
upper member
10,000
M31
A/AC/C
fSL
254
19.89
0.08
Modesto,
upper member
10,000
M46
A/AC/Cox
fSL
250
33.73
0.13
Modesto,
lower member
20,000 to
70,000
M12
A/BI/B3/C
SL
413
67.05
0.16
Riverbank,
upper member
130,000
R9
A/B/B3/Cox
SL/LS
400
115.85
0.29
Riverbank,
upper member
130,000
R33
A/Bl/B3/Cox
SL/LS
300
87.85
0.29
Turlock Lake
600,000
T6
A/Bt/BC
SL
190
148.10
0.30
Turlock Lake
600,000
Til
A/Bt/BC/Cox
SL
500
148.78
0.78
Note: data from Harden and Marchand, 1977; Marchand and Allwardt, 1981; Harden, 1982; Harden and Taylor, 1983; and Harden, 1986. Profile index values
determined on the basis of eight properties.
*f, fine; LS, loamy sand; SL, sandy loam; SiL, silty loam.
Clay and Fe2C>3d content are correlated
(Spearman rank correlation, linear regression),
especially in latest Pleistocene and the late Pleistocene soils (minimum a = 0.01, range in value of
r = 0.78 to 0.98). This relation has been observed in soils formed on fluvial deposits elsewhere in the Transverse Ranges (McFadden and
Hendricks, 1985) and suggests that processes resulting in the accumulation of these two components in soils are genetically related. Statistical
analysis shows that silt and Fe2C>3d contents are
not correlated or are weakly correlated in latest
and late Pleistocene soils, in contrast to Holocene soils. Furthermore, silt and FeOT contents
are more poorly correlated in latest Pleistocene
soils than in Holocene soils and are actually
weakly negatively correlated in the soil on the
late Pleistocene terrace (Qoa-d) (r = -0.66).
These results indicate that iron oxide and clay
accumulation are both increasingly related to
chemical weathering rather than to incorporation of eolian dust. As shown previously, alteration of ferrous iron in Pleistocene soils creates
authigenic ferric-iron oxides. Oxidation of ferrous iron in the eolian silt must account for
weak or even ultimately negative correlation of
FeOT and silt contents because these components are so strongly correlated in young unweathered soils. Authigenic clay results from
hydrolytic weathering of most minerals; hence,
the strong correlation of clay and Fe203d content in Pleistocene soils is at least partly due to
co-formation of authigenic clay and ferric-iron
oxides. These conclusions agree with regional
studies of the clay mineralogy that identify kaolinite as the ultimately predominant mineral in
late and middle Pleistocene soils of the Transverse Ranges (McFadden and Hendricks, 1982).
Because little kaolinite is present in soils
throughout the Mojave Desert (McFadden,
1982; McFadden and others, 1986; McFadden
and Bull, 1987), eolian dust derived from this
region can probably supply little kaolinite to terraces and fans of the Transverse Ranges, implying a primarily authigenic origin for kaolinite in
the Cajon Pass area.
RECOGNITION OF A PEDOLOGIC
THRESHOLD AND IMPACT ON
RATES OF SOIL DEVELOPMENT
The observed contrasts in rates of soil development with respect to aspects of A- and Bhorizon development are the result of the
interdependent nature of factors that influence
soil development through time (Yaalon, 1971;
Jenny, 1980; Birkeland, 1984). The initially
rapid rate of A-horizon development in the
study area, for example, reflects the initially
permeable nature of a parent material that favors incorporation of eolian dust, a sufficiently
moist climate, and a vegetation cover that provides abundant soil organic matter. Despite low
pH and intense winter leaching, chemical alteration in these soils is limited to very slight alteration of ferrous iron, weak soil reddening, and
partial grussification of large stones. Accumulated soil materials apparently consist almost entirely of material derived from eolian dust. As
the A horizon develops, soil permeability, waterholding capacity, and infiltration change, eventually favoring accumulation of clay. Accumula-
291
tion of clay, silt, and organic matter creates
positive feedback that accelerates changes in soil
permeability and infiltration rates. Subsequent
soil development is progressively characterized
by increasing accumulations of authigenic ferriciron oxide and clay due to chemical weathering
under conditions of acidic pH and strong leaching. The transition from a permeable, noncolloidal soil environment favoring dust incorporation
to a more strongly colloidal, less permeable system favoring chemical weathering occurs over a
relatively short period of time and constitutes an
extrinsic pedologic threshold. Soil thresholds,
extrinsic or otherwise, have been recognized or
suggested in previous studies (for example,
McFadden, 1981; Muhs, 1984; Birkeland,
1984) and are analogous to the geomorphic
thresholds described by Schumm (1973, 1977,
1979) and Coates and Vitek (1980); in each
situation, relatively constant processes produce
sharp changes in the rate of formation of soil
properties or analogous geomorphic parameters.
Comparison of morphologic data from the
chronosequence of soils in the Merced area in
the San Joaquin Valley of California (Table 3)
(Harden and Marchand, 1977; Marchand and
Allwardt, 1981; Harden, 1982; Harden and Taylor, 1983; Harden, 1986) with data for the
Cajon Creek chronosequence (Tables 1 and 2)
permits an evaluation of the variables that may
influence the timing or relative significance of a
pedologic threshold. The present climate of the
Merced area (Mediterranean; mean annual precipitation = 410 mm, mean annual temperature
= 16 °C) is quite similar to the climate in the
Cajon Creek area. The parent materials in the
Merced area, however, are typically finer
grained, consisting typically of sandy loam, and
are chiefly granitic. Deposit or surface ages are
based on a variety of data, including 14C, uranium trend, K-Ar dating methods, and correlations with the marine oxygen-isotope stages.
Values of the Merced profile index also were
calculated for soil depths that significantly exceed depths for which the index was calculated
for most Cajon Creek soils. Differences in the
index value due to thickness, however, can
be significantly reduced by determining the
weighted mean value of the index (Harden and
Taylor, 1983).
Comparison of the two study areas indicates
many similarities with respect to the morphological trends in soil development. The soils on the
upper member of the Modesto Formation
(-10,000 yr old, Harden, 1986), however, apparently are more weakly developed than are
the soils on latest Pleistocene terraces of the
Cajon Creek area and are considerably more
292
McFADDEN A N D WELDON
similar to those on the middle and early Holo- dust. Reheis observed logarithmic rates of soil
cene terraces. The youngest deposit on which an development in more humid regions of Wyoargillic-horizon-bearing soil is present is the ming, inferred to reflect the increased signifi20,000- to 70,000-yr-old lower member of the cance of chemical weathering. As noted preModesto Formation (Marchand and Allwardt, viously, long-term rates of soil development
1981; Harden and Taylor, 1983; Harden, 1986). were logarithmic in the Cajon Pass region. If
The timing of argillic horizon development in only soils of middle Holocene age or younger
the Merced area is therefore slightly to much are considered, however, some rates of soil deolder than the timing determined for the Cajon velopment in the Cajon Pass area can be deCreek area. Note that values of the weighted scribed as linear: profile Fe2C>3d = 0.002 (terrace
means of soils on the 130,000-yr-old upper unit age) + 4.84. An initially linear rate of developof the Riverbank Formation are almost identical ment may reflect the predominance of dust into those of the weighted means of the soil on the corporation. The over-all logarithmic rate re55,000-yr-old Qoa-d deposit. The 130,000-yr- flects the processes of dust incorporation and an
old age assignment, however, has been deter- increasing magnitude of chemical weathering,
mined only on the basis of uraniun-trend dating and it masks the early linear-rate phase as well
and correlation to the oxygen-isotope record as the thresholds that occur and are discernable
and therefore is subject to some uncertainty. If over only relatively short periods of time.
that age is correct, an over-all slower rate of
morphological development in the Merced area
CONCLUSIONS
compared to that in the Cajon Creek area is
indicated, at least during the initial 130,000 yr of
Studies of the well-dated sequence of soils in
soil development. The 14 C dates within deposits
the Cajon Pass area demonstrate that many soil
of the upper member of the Modesto Formation
characteristics change systematically with time.
demonstrate a latest Pleistocene age, which is
The rates and magnitude of soil development
consistent with the outwash origin attributed to
and the particular processes dominating soil dethis deposit by Marchand and Allwardt (1981).
velopment, however, have varied significantly
Because the lower member of the Modesto is
through time. The most important single variaolder than latest Pleistocene, the threshold of soil
ble affecting the initial phase of soil development
development that was recognized in our study
at Cajon Pass is the incorporation of eolian dust,
probably was crossed in the Merced area after a
which is the primary source of silt, most secondmuch longer period of soil development than is
ary iron oxides, and some clay. The continuing
required in the Cajon Creek area. Differences in
accumulation of these components subsequently
soil parent materials, eolian influx rates, or other
changes the initially permeable, noncolloidal soil
as yet unknown factors account for the different
environment to an increasingly less permeable
times required to attain the threshold in the two
and more colloidal environment, which in conareas. The great difference in initial permeability
junction with strong seasonal leaching and
of the fine-grainei Merced area soils compared
acidic pH, promotes an increasing degree of
to that in Cajon Pass probably in particular rechemical weathering of the soil parent materials
duces the significance of the threshold.
and the incorporated aerosolic materials. MoreIn an environment more arid than that of over, steady-state conditions of soil developCajon Pass, the threshold identified in this study ment are not attained over a time span of at least
may be difficult to recognize. Shallow leaching half a million years, a conclusion also drawn by
and a relatively high rate of dust influx min- Muhs (1982) and by Harden and Marchand
imizes the rate and magnitude of chemical (1977) in studies of soils elsewhere in California.
weathering relative to the rate and magnitude of
soil plasma accumulation by incorporation of
dust, and they generally produce much lower
rates of soil development than those observed in
the study area and elsewhere in the Transverse
Ranges during a period of several hundred thousand years (McFadden, 1982; McFadden and
Bull, 1987). Reheis (1984), for example, reported slow linear rates of soil development in
arid regions of Wyoming, presumably attributable to the predominance of incorporation of
The pedologic threshold is recorded in the
Cajon Pass area by the rapid appearance of the
argillic horizon in soils that are only -3,000 yr
older than those having only color B horizons.
In Cajon Pass, the threshold requires -8,000 yr
to occur, but the timing and relative importance
of this threshold in other regions are presumably
affected by several variables, such as dust influx,
soil-water balance, and initial parent-material
characteristics. In arid environments, dust incorporation may always dominate soil develop-
ment, whereas in climates more humid than that
of the study area, intense leaching and chemical
alteration may always dominate soil development. Regions that have significant dust influx
rates a nd moderately intense leaching, such as
the Cajon Pass area, may be unique in that periods of time can be identified during which one
process dominates over the other. Implicit in the
concept of such a threshold is that episodes of
rapid rates of soil development do not necessarily require climatic regimes that were uniquely
favorable for chemical weathering. The threshold could be influenced significantly, however,
by climatic changes that cause changes in rate,
magnitude, and composition of aerosolic dust
influx. Changes in soil properties in a given sequence of soils may thus be quite systematic, but
significant differences in the degree of soil development on late Pleistocene and Holocene deposits that were observed in different regions may
occur, despite similar soil ages and climates, due
to contrasting geomorphic settings that influence
the local timing and relative importance of a
pedologic threshold. An important implication
of this hypothesis is that caution should be exercised in the assignment of ages or correlations
among late Quaternary deposits on the basis of
comparison of soils that exhibit relatively similar
pedologic characteristics.
ACKNOWLEDGMENTS
The authors thank C. Prentice, G. Martinez,
L. Smith, P. Karas, M. Jackson, and T. Bullard
who provided valuable assistance during the
field and laboratory phases of this study. Reviews of a preliminary manuscript of this paper
by J. C. Tinsley, J. W. Harden, and V. T. Holliday resulted in improvements in this paper.
Laboratory analysis and field work conducted
during this study were partly supported through
grants (U.S. Geological Survey Hazards Reduction Programs, Contract Numbers 14-08-000116774, 18285, 19756, and 21275) to Kerry E.
Sieh, who served as thesis adviser to R. Weldon
during this project.
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1986
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