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TWO
Climate, Paleoclimate, and Paleovegetation
R I C HAR D A. M I N N I C H
I NTRODUCTION
MODE R N CLI MATE
Temperature
Winter Precipitation
The North American Monsoon
Potential Evapotranspiration and Runoff
Climatic Variability in the Instrumented Record
El Niño-Southern Oscillantion
Pacific Decadal Oscillation
TE RTIARY CLI MATE CHANG E AN D DEVE LOP M E NT
OF CA LI FOR N IAN VEG ETATION
Cenozoic Global Cooling
Plate Tectonics and Topography of the
Western United States
Paleovegetation
Late Cretaceous to Early Eocene
Middle Eocene and Oligocene
Miocene to Early Pleistocene
Late Pleistocene and Holocene
Late Pleistocene
Early and Middle Holocene
Mid-Holocene
Late-Holocene drought
Little Ice Age
S U M MARY OF TE RTIARY–QUARTE R NARY
CLI MATE AN D VEG ETATION CHANG E
AR EAS FOR F UTU R E R E S EARCH
Introduction
The extraordinary variety of climate in California—ranging
from high rainfall regimes in the northwestern mountains to
Death Valley, the driest place in North America, is accompa-
nied by perhaps the greatest regional species diversity in temperate North America. The vegetation has comparable richness
in life forms that includes tall conifer forests, riparian deciduous hardwoods, broadleaved evergreen woodlands and savannas, shrublands of evergreen chaparral and drought-deciduous
shrubs, and fields of annual and perennial forbs and grasses, all
of which cover extensive areas in the state. The vegetation is
an overprint of floras of different age and regional origins that
are vestiges of Tertiary vegetation and climate history. This
chapter first summarizes the modern climate of California for
background to the reviews of vegetation in this volume. What
follows is a summary of paleoclimate and paleobiogeography
of the California flora since the late Cretaceous.
Modern Climate
California’s mediterranean climate of winter precipitation
and protracted summer drought is an outcome of seasonal
changes in global circulation. In winter, strong latitudinal
temperature gradients focus the polar-front jet stream and
onshore flow of moist surface westerlies around surface low
pressure in the Gulf of Alaska into northern California and
the Pacific Northwest. In summer, reduced latitudinal temperature gradients weaken the jet stream that shifts poleward to western Canada. Gulf of Alaska low pressure is
replaced by surface high pressure covering the northeast
Pacific Ocean. A vital component of the mediterranean climate is the coastal marine layer, a steady-state feature associated with the cooling and moistening of the tropospheric
boundary layer overlying the cold, upwelling California
Current. It is capped by a strong thermal inversion that
divides the layer from warm and dry subsiding air masses
aloft. The marine layer virtually precludes convective precipitation of Pacific Ocean air masses from May to September. Summary climate data are given in Table 2.1.
43
123 26
123 15
122 24
121 54
120 28
118 24
117 10
39 31
38 31
37 37
36 36
34 55
33 56
32 44
Fort Bragg
Fort Ross
San Francisco AP
Monterey
Santa Maria
Los Angeles Int’l AP
San Diego Int’l AP
119 43
119 03
117 23
36 47
35 26
33 57
Fresno Air Terminal
Bakersfield
Riverside
120 49
120 57
121 00
38 42
39 07
39 15
Placerville
Colfax
Nevada City
N. Sierra Nevada
121 30
38 30
Sacramento AP
Orland
122 12
122 25
40 43
Shasta Dam
39 45
121 59
40 59
Pit River
Cen. Val./ Inter. SCA
123 58
41 51
Gasquet
Siskiyou/ Cascades
124 10
40 49
Eureka
Pacific Coast
Station/Region
Lat./Long.
(deg and min)
847
731
563
275
149
0.2
0.3
0.2
0.1
0.0
0.0
0.1
4
109
0.2
0.5
0.9
0.9
0.1
0.1
0.1
0.2
0.1
0.3
0.3
0.4
Jul
77
327
445
117
4
30
79
116
2
33
35
6
Alt.
(m)
0.3
0.5
0.2
0.3
0.1
0.1
0.2
0.4
1.2
1.5
2.6
0.2
0.2
0.1
0.3
0.1
0.7
0.9
0.9
Aug
TA B L E 2.1
1.8
2.0
1.5
0.7
0.3
0.5
0.7
1.0
3.4
5.3
5.8
0.5
0.4
0.6
0.7
0.5
1.3
1.6
2.2
Sep
7.0
5.8
5.5
0.7
0.7
1.3
2.3
2.8
7.7
11.6
15.4
1.2
0.8
1.2
2.4
2.1
6.3
6.6
6.9
Oct
16.3
16.7
11.6
2.4
1.5
3.0
5.4
7.2
21.0
28.7
37.9
2.5
3.8
3.4
6.0
6.1
13.1
13.7
14.6
Nov
23.0
20.3
16.1
3.1
1.9
3.8
7.1
8.2
25.5
30.5
40.4
4.5
4.3
4.6
7.5
8.6
15.6
17.4
16.4
Dec
26.0
22.8
18.1
5.8
2.7
4.9
9.5
11.0
31.2
33.1
37.7
5.3
7.0
6.4
10.9
11.7
20.9
19.8
17.3
Jan
24.8
19.4
17.2
4.7
2.9
4.9
7.9
8.8
25.9
27.1
30.7
4.9
6.6
7.1
8.4
9.1
15.5
16.2
13.5
Feb
Mean Precipitation (cm)
20.0
18.3
14.7
4.5
3.0
4.9
6.0
7.5
22.9
27.6
32.6
4.4
5.1
6.0
8.2
8.4
13.5
14.8
13.5
Mar
Climatological Data of Selected Stations in California
10.9
9.1
7.8
2.0
1.7
2.5
2.8
3.1
10.7
13.6
16.3
2.0
2.0
2.7
3.9
3.3
6.7
7.3
7.2
Apr
5.3
4.3
3.9
0.6
0.5
0.9
1.2
1.9
6.1
6.4
9.6
0.5
0.4
0.7
1.2
1.0
2.4
3.5
4.3
May
1.6
1.5
1.2
0.2
0.2
0.4
0.4
1.0
3.0
2.3
3.6
0.2
0.2
0.1
0.5
0.3
1.1
1.1
1.6
Jun
137.2
120.6
97.3
25.1
15.5
27.2
43.6
53.1
159.1
188.6
233.5
24.3
30.9
32.6
50.2
50.9
97.3
103.2
98.8
Total
50.8
36.1
7.6
0
0.2
0.2
0
1.0
12.7
–
0
0
0
0
0.1
0
0
0
0.5
1
0
0
0
0
0
0
0
13
–
0
0
0
0
0
0
0
0
0
Days Mean
2.5cm
Snowfall
(cm)
4.5
6.9
5.8
12.1
8.8
7.7
7.5
7.2
7.4
–
–
13.4
13.3
10.6
10.9
9.4
9.5
8.7
8.8
Jan
21.3
24.6
23.4
25.7
28.7
27.5
24.1
25.8
27.5
–
–
20.8
20.6
17.1
15.5
17.0
13.9
13.9
13.9
Jul
Mean
Temp (EC)
–
225
170
257
287
303
321
273
–
–
–
365
365
272
365
365
–
293
335
Frost2
free
days
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119 56
120 11
120 08
38 36
39 20
39 10
Tamaracka
Truckee RS
Tahoe City
119 47
118 58
37 57
36 44
36 36
37 56
Hetch Hetchy
Grant Grove
Lodgepolea
a
34 18
33 45
Mt. Wilson
Idyllwild
33 48
Mt. San Jacinto
41 30
39 58
38 15
37 35
Alturas
Doyle
Bridgeport
White Mountain
Modoc Plat./Mono L.
34 15
Big Bear Lake
34 15
34 15
Lake Arrowhead
Big Bear Lake Dam
33 14
Henshaw Dam
SCA Mountains
Ellery Lake
118 52
36 28
Three Rivers
118 14
119 14
120 05
120 33
116 38
116 53
116 59
116 42
118 04
117 11
116 46
119 14
118 44
119 02
36 23
Lemon Cove
b
120 38
39 19
Lake Spaulding
S. Sierra Nevada
121 06
39 34
Strawberry Valley
3801
1972
1338
1341
2566
2070
2073
1637
1740
1586
823
2926
2051
2011
1179
346
156
1898
1834
2457
1571
1174
2.7
1.2
0.8
0.8
1.4
1.9
1.1
1.8
0.2
0.4
0.8
1.5
1.1
0.6
0.5
0.2
0.0
0.7
0.9
1.0
0.5
0.4
2.8
1.2
0.7
0.9
4.1
2.5
1.9
2.1
0.5
0.9
1.3
1.7
0.8
0.3
0.5
0.2
0.1
0.8
1.2
0.5
0.8
0.8
2.2
1.2
1.1
1.2
3.3
1.4
1.9
2.2
1.9
2.3
1.3
1.3
3.7
3.1
1.8
1.5
0.7
1.6
1.9
1.8
2.8
3.3
2.6
0.7
1.7
2.4
1.7
1.8
3.5
2.4
2.8
3.8
2.0
3.6
5.0
5.2
5.3
2.8
1.7
4.6
4.2
7.4
10.0
11.6
3.1
2.3
3.5
3.7
9.4
5.5
7.7
6.0
9.7
10.7
6.1
5.1
11.3
11.6
10.6
6.9
4.1
9.3
10.0
10.7
20.9
26.5
6.6
2.7
4.6
3.8
12.0
7.5
13.4
8.7
11.3
14.1
8.1
11.0
16.0
14.7
14.3
7.8
5.0
13.7
13.5
18.9
28.6
35.0
6.3
3.6
5.3
3.9
12.0
11.0
17.9
13.1
19.3
22.3
13.1
11.2
25.7
20.3
15.4
11.9
6.9
15.7
16.2
23.0
31.2
42.0
4.3
3.8
3.6
3.6
9.2
10.6
18.1
11.3
19.6
20.2
11.9
11.2
23.0
17.9
14.7
10.7
5.9
14.0
12.6
28.0
29.0
33.8
5.5
2.4
3.1
3.5
7.4
8.7
16.3
10.9
16.0
17.3
12.3
8.0
19.1
19.1
13.3
11.8
6.3
10.4
10.7
21.3
24.9
30.0
4.8
1.0
1.6
2.7
2.7
3.4
6.1
5.0
7.0
8.0
5.1
4.9
7.8
7.7
8.2
4.6
3.3
5.4
5.5
5.6
14.1
14.9
5.0
1.3
1.9
3.4
1.4
1.3
2.5
2.0
2.1
3.4
1.7
2.6
3.2
3.7
4.6
2.0
1.3
3.0
3.5
4.2
8.3
7.8
2.1
1.4
1.4
2.5
0.2
0.4
0.4
0.4
0.5
0.5
0.3
2.1
1.8
3.7
2.1
0.8
0.4
1.8
1.8
3.3
3.0
2.4
48.0
22.8
29.0
33.4
64.8
56.0
91.4
65.9
90.9
103.9
64.0
64.2
118.5
107.8
91.3
61.2
35.7
81.0
82.0
125.7
174.1
208.5
402.6
106.9
57.9
82.2
–
161.5
313.1
100.1
68.6
123.6
3.8
664.5
644.8
493.5
165.6
0
0
479.8
527.5
1098.8
650.2
277.6
0
–
41
22
21
–
–
–
17
–
14
–
167
147
33
0
0
131
139
–
140
51
16.4
7.4
3.4
9.3
–
68
–
–
–
–
–
–
–
–
–
–
–
121
180
–
279
–
42
–
101
–
(continued)
22.6
18.8
15.9
0.8
1.6
0.0
17.5
16.3
1.9
0.9
20.1
23.6
20.5
22.3
12.7
15.1
17.2
21.6
27.8
27.1
16.1
16.4
13.4
17.8
20.1
4.5
6.6
2.9
6.9
-4.8
-2.7
1.2
3.5
8.5
8.0
-1.8
-2.8
-3.8
1.2
3.9
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35 03
34 51
34 46
36 28
Palmdale
Mojave
Daggett-Barstow AP
Needles
32 51
34 08
33 37
32 53
Thermal AP
Imperial
Iron Mountain
Blythe
114 52
114 36
115 08
115 34
116 10
116 02
115 33
115 33
117 11
116 52
116 47
116 48
118 10
118 06
118 21
0.3
19
147
81
0.7
0.7
0.7
0.5
34
281
1.5
2.7
2.0
0.9
0.3
0.9
1.0
0.3
0.1
0.4
Jul
602
1441
1325
1249
-59
278
585
832
791
1250
Alt.
(m)
1.4
1.7
1.0
0.7
0.7
1.8
3.2
3.8
1.6
0.4
1.7
1.0
0.5
0.5
0.3
Aug
0.9
1.0
0.7
0.7
0.7
1.2
1.6
2.1
1.2
0.5
1.2
0.9
0.6
0.5
0.5
Sep
0.9
0.7
0.9
0.6
0.4
0.6
1.2
1.6
0.7
0.3
0.8
0.4
0.5
0.7
0.5
Oct
0.7
0.5
0.6
0.5
0.7
0.6
1.6
1.8
1.1
0.5
0.9
0.6
1.6
1.8
1.3
Nov
1.3
1.1
1.1
0.9
0.8
1.0
1.5
2.4
1.4
0.4
1.1
1.1
1.9
3.5
2.1
Dec
1.6
1.2
1.4
1.1
1.4
1.3
2.5
3.2
2.4
0.7
1.6
1.6
3.0
3.9
2.8
Jan
1.1
1.0
0.8
0.8
1.2
0.9
2.1
3.9
3.1
1.2
1.2
1.1
2.9
3.9
2.4
Feb
Mean Precipitation (cm)
0.9
0.9
1.0
0.7
0.9
0.9
2.3
3.6
2.8
0.9
0.9
1.1
2.3
3.3
1.3
Mar
0.3
0.5
0.4
0.2
0.2
0.3
1.1
1.4
0.8
0.3
0.3
0.6
0.8
1.2
0.8
Apr
0.2
0.1
0.2
0.1
0.2
0.3
0.8
0.7
1.1
0.2
0.2
0.2
0.3
0.3
0.7
May
0.0
0.1
0.1
0.0
0.0
0.0
0.6
0.3
0.4
0.1
0.1
0.2
0.1
0.1
0.4
Jun
10.0
9.5
8.9
6.6
7.6
10.4
21.2
25.8
17.5
5.8
10.9
9.8
14.8
19.8
13.5
Total
0
0
0
0
0
2.3
23.3
8.6
6.0
0
0.8
2.0
3.8
3.8
20.5
0
0
0
0
0
0
6
0
0
0
0
0
0
0
5
Days Mean
2.5cm
Snowfall
(cm)
13.8
12.2
12.0
13.1
12.6
9.6
–
7.8
4.9
11.4
11.5
9.0
7.4
7.3
2.9
Jan
34.1
34.8
34.8
33.1
32.9
31.3
–
28.0
26.3
38.3
35.4
31.6
28.1
27.3
24.8
Jul
Mean
Temp (EC)
–
290
–
317
317
266
–
–
–
–
316
307
–
–
–
Frost2
free
days
Western Regional Climate Center, Desert Research Institute. Station record 1948–2001. Web site: www.wrcc.dri.edu/index.html
Base OEC. Source: U.S. Department of Commerce. Weather Bureau. Climates of the States, California.
b
Department of Commerce. Weather Bureau. Climatic summary of the United States, Supplement for 1931–1952.
c
Climatological Data, California. Monthly summaries. National Oceanic and Atmospheric Administration. Environmental Data and Information Service. National Climatic Center. Asheville, North Carolina.
a
NOTE :
Gold Rock Ranch
33 38
Twenty-nine Palms
Sonoran Desert
34 08
35 28
Mitchells Cavern
Mountain Pass
34.57
Wildrose RS
Mojave Desert Mtns
36 16
33 35
Bishop AP
Death Valley
37 22
Mojave Desert
Station/Region
Lat./Long.
(deg and min)
TA B L E 2.1 (continued)
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Temperature
The position of the jet stream is coupled with broadscale latitudinal temperature gradients in California. In January
mean sea-level temperatures range from 9C in northern
California to 14C in southern California. Atmospheric
lapse rates result in temperatures decreasing to 0C at ca.
1,600 m in northern California and 2,200 m in southern
California, with means as low as –6C along the crest of the
Sierra Nevada. Many low-lying basins contain persistent
temperature inversion layers resulting from radiational
cooling and low insolation. In the Central Valley, where
ground inversions are maintained by reflective ground fogs,
mean January temperature averages 8C. Surface inversions
result in means as low as –3C in the high northeastern
plateaus from Modoc to Lake Tahoe.
From May to September, California is dominated by strong
onshore flows from the Pacific anticyclone to thermal low
pressure over the hot desert interior. Northwesterly gradient
winds combined with sea breezes and anabatic circulations
transport the marine layer air inland to the coastal ranges
usually within 100 km of the coast. Farther inland, the
marine layer dissipates from diabatic heating and mixing
with warm air aloft (Glendening, Ulrickson, and Businger
1986). Mean July temperatures along the coast reflect local
sea surface temperatures, ranging from 14C–16C in strong
upwelling zones north of Point Conception to 18C–22C in
the southern California bight. Temperatures increase to
24C–28C in the inland valleys of southern California and
along the Central Valley. The deserts beyond the reach of the
marine layer average 26C–35C, and 38C in Death Valley.
Temperatures in mountains reflect ambient lapse rates,
decreasing to 14C–18C at 2,500 m and 10C at 3,000 m.
Mountain and desert temperatures and lapse rates above the
marine layer are isoclinal with latitude across California from
June to September. The frost-free period varies primarily with
elevation and distance from the Pacific Ocean. Freezing temperatures are rare along the Pacific coast south of San Francisco and infrequent in December and January as far north as
the Oregon border. Mild freezes occur from November to February in the Central Valley, interior valleys of southern California, and the Salton Sea trough. The frost-free period is less
than 100 days above 1,500–2,000 m in the Modoc Plateau,
Sierra Nevada, and mountains of southern California.
Winter Precipitation
Winter precipitation results primarily from cold fronts of
extratropical cyclones and associated troughs of the jet
stream moving into the region from the North Pacific Ocean.
Because the mean position of the jet lies from the Pacific
Northwest to northern California, the frequency of storms
decreases southward in the state. The deepening of the
marine layer before the passage of cold fronts gives rise to
weather conditions similar to frontal occlusions, with extensive cloud shields and long periods of steady precipitation in
stable air. Winds aloft are southwesterly because cold fronts
precede the passage of the associated trough aloft. Low-level
winds are south to southeasterly (Minnich 1984). In the postfrontal phase, steady precipitation is replaced by convective
showers concentrated over high terrain, with the winds at
the surface and aloft shifting to westerly. Clouds and showers
dissipate with the onset of subsidence following the passage
of troughs, with winds turning northwesterly.
The interaction between prefrontal circulation and terrain results in strong gradients in mean annual precipitation throughout California. Because storm air masses are
stable, the variation in local precipitation is more influenced by mechanical (physiographic) lift over mountain
barriers rather than from thermal convection. Physiographic lift is most intense on the south- to southwest-facing
escarpments that lie at right angles to storm winds. Amounts
decrease downwind to inland mountain ranges—regardless
of altitude—due to depletion of storm air mass moisture and
descending airflow in rain shadows.
The average annual precipitation (AAP) along the northern California coast varies from 50 cm at San Francisco to
100 cm north of Ft. Ross, with locally higher amounts
where mountains skirt the coastline. The coastal mountains
north of Eureka and near Cape Mendocino receive 250 cm.
Totals in the North Coast Ranges decrease downwind to
150–200 cm in the Salmon and Siskiyou Ranges, and
decrease southward with the declining general altitude of
the mountains to 100 cm near Santa Rosa. Rain shadows
from the North Coast Ranges decrease AAPs to 35–60 cm in
the Sacramento Valley. However, low-level convergence of
southerly prefrontal air masses against the narrowing rift
between the Sierra Nevada and North Coast Range increases
AAP to 60–80 cm north of Red Bluff. Orographic lift along
the undissected western slope of the northern Sierra Nevada
increases AAP from 60 cm along the lower foothills to
150–200 cm at the crest. To the east, rain shadows result in
AAPs of 20–60 cm in the Modoc Plateau, and 60–80 cm in
the Lake Tahoe Basin.
In the South Coast Ranges, the AAP is 100–150 cm on the
steep coastal escarpments of the “lead” Santa Cruz and
Santa Lucia Mountains but amounts in the “downwind”
Diablo Ranges seldom exceed 50 cm. Intervening basins
including Salinas Valley receive 30–40 cm. Rain shadows
extending from the South Coast Ranges into the San
Joaquin Valley produce an AAP of 30 cm near the Sacramento delta, lowering to 15 cm near Bakersfield and the
Carrizo Plain. Amounts then increase with orographic lift to
30–50 cm in the Sierra Nevada foothills. The topographic
complexity of the southern Sierra Nevada results in large
variability of AAP along the west slope. Steep southwestern
exposures have AAP of 100–150 cm, as at Yosemite, the
upper San Joaquin drainage, Kaiser Ridge, and from Sequoia
National Park to the Great Western Divide and Greenhorn
Mountains. Leeward slopes on the coastal front receive
50–100 cm, including the upper Tuolumne River, Mono
Creek Basin, the upper Kings River, and the Kern River
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plateau northward to Mt. Whitney. The AAP seldom
exceeds 50 cm south of the Greenhorn Mountains due to
low altitude of these ranges and their leeward position to
the western Transverse and South Coast Ranges.
In southern California, precipitation is highest on the
steep southern escarpments of the Transverse Ranges (AAP,
80–110 cm). Amounts decrease to 60–80 cm in the “downwind” San Rafael Mountains, Pine Mountain Ridge, and leeward slopes of the San Gabriel and San Bernardino Mountains. Farther inland, the relatively high ranges that include
Mt. Pinos, San Emigdio, Tehachapi and Liebre Mountains,
and drainages north and east of Big Bear Basin receive 35–60
cm. The AAP is only 40–60 cm on the coastal slopes of the
Peninsular Ranges because storm winds frequently parallel
the escarpment. Amounts reach 80–100 cm on local southern escarpments of the Santa Ana Mountains, Palomar
Mountain, and Cuyamaca Peak. The high San Jacinto
Mountains receive only 40–70 cm and the Santa Rosa
Mountains only 40–50 cm due to their leeward position to
the Santa Ana and Palomar Mountains. Amounts in the
southern California plains vary from 25–35 cm at the coast
to 40–50 cm at the base of the mountains.
The AAP in the southeastern deserts and Owens Valley is
mostly 10–15 cm, with totals of 6–10 cm in the Salton Sea
trough and 5.8 cm in Death Valley. Amounts reach 30 cm in
the Panamint and Inyo Mountains, and the higher ranges
in the northeast Mojave Desert. The White Mountains
above 3,000 m receive 50 cm.
The peak of the winter precipitation season shifts from
December/January in northern California to January/February
in the south. This trend reflects a gradual equatorward shift
of the jet stream due to the gradual cooling of the North
Pacific SSTs through the winter. The precipitation maximum reverts to December/January in the southeastern
Deserts when rare cyclones or cutoff lows with southerly
wind trajectories advect moisture from tropical Pacific and
along the Gulf of California. Precipitation declines rapidly
throughout the state in April due to the weakening of the
jet stream. Weak disturbances extend the rainy season into
May and June in the northern Sierras, the Modoc Plateau
and other interior basins south to the White Mountains
when relatively warm Pacific air masses compared to winter
reduce the Sierra Nevada rain shadow, resulting in convective showers on leeward slopes under high sun. Extratropical precipitation virtually ceases throughout the state by
late June when the jet stream assumes its summer position
near the U.S.–Canadian border.
An accumulating winter snowpack melt serves to increase
dry season soil moisture in high mountain watersheds. The
ratio of the water content of frozen precipitation to the
average annual precipitation (S/AAP) increases uniformly
with altitude reflecting atmospheric lapse rates (Minnich
1986). The lower limit of reliable snowfall of 1,000–1,200 m
in California approximates the moist adiabatic lapse rate
from mean sea-surface temperatures to storm freezing lines.
In southern California, the S/AAP ratios increase to 25% at
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1,750 m, 75% at 2,750 m, and 100% at 3,000 m. Average
snow lines are about 200–400 m lower in the Sierra Nevada
(Barbour et al. 1991), and ca. 400–600 m lower in northern
California. The water content of solid precipitation depends
on both the S/AAP ratio and on the total AAP. In general,
frozen precipitation exceeds 50 cm above 2,300 m in southern California, 1,900 m in the central Sierra Nevada, and
1,400 m at Mt. Lassen. Amounts exceed 100 cm above 3,000
m in southern California, 2,500 m in the Sierra Nevada, and
2,000 m in northern California.
S/AAP ratios decline through the rainy season. In southern California, average snow lines decrease from 2,700 m in
November to 2,300 m in February and 1,700 m in March
and April (Minnich 1986). Seasonal trends have not been
documented for northern California. Late fall and winter
storms advect moisture from the subtropical Pacific,
whereas storms in March and April are strongly influenced
by the marine layer overlying the ever-cooling California
current. Average snow levels in California tend to rise
with increasing total annual precipitation due largely to
the advection of moist subtropical storm air masses during
El Niño events. In southern California, average annual snow
lines vary from 2,000 m in years with 70% normal precipitation to 2,400 m in years with 140% of normal. Extraordinary snow accumulations during very wet years are frequently limited to the highest elevations, with middle
elevations producing storm runoff. The average snowline
during the southern California floods of January 1969 was
2,700–3,000 m. The snowline during the 1998 New Years
flood at Yosemite was 2,500 m.
Snowmelt is dependent primarily on solar and infrared
radiant loading (Miller 1981). In winter, melt rates exceed
snow accumulation rates on south-facing exposures lying
close to right angles to the sun below ca. 2,000 m in southern California (Minnich 1984) and ca. 1,300 to 1,700 m in
the Sierra Nevada and North Coast Ranges. With increasing
elevation, snowmelt is less sensitive to slope aspect because
the upslope snowpack retreat is phased with ever-higher
sun angles in spring and summer.
The North American Monsoon
From July to September, the western margin of the North
American monsoon, a deep layer of moist, unstable tropical
air periodically causes afternoon thunderstorms in California, especially in the eastern mountains and deserts. The
monsoon moves into the region around a mid-tropospheric
anticyclone sustained by intense convective heating of high
elevation land surfaces of the southwestern United States
and Mexican plateau (Hales 1974). Tropical moisture of the
monsoon arrives from the equatorial eastern Pacific Ocean
by way of the Gulf of California. When the anticyclone center lies over northern Mexico, monsoon moisture is steered
northeastward from the Gulf of California into northwestern
Mexico and Arizona. Dry southwesterly flow over California
results in clear skies, although deep troughs in the westerlies
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produce infrequent thunderstorms that source Pacific air
masses in the Cascades/Siskiyou Mountains and far northern
Sierra Nevada. When the anticyclone moves northward into
the southwestern United States, southeasterly winds aloft
transport monsoon moisture into southwestern California
and the Sierra Nevada. Moisture arrives at upper levels (3–4
km) as convective debris from thunderstorms or mesoscale
convective systems over the Sierra Madre Occidental of
northwestern Mexico and Gulf of California. Below 2 km,
moist air masses (derived from convective outflows of thunderstorms over Mexico) surge into the Salton Sea trough,
Colorado River Valley, and occasionally as far north as
Owens Valley (Hales 1974; Stensrud, Gall, and Nordquist
1997). Moisture surges also result from the lift of the trade
wind layer overlying the Gulf of California by passing east
Pacific tropical cyclones.
Convection is most frequent in the mountains exposed to
low-level moist air masses from the Gulf of California, primarily east of a line from the Peninsular Ranges and eastern
San Bernardino Mountains to the eastern escarpment of the
Sierra Nevada northward to Lake Tahoe. Total July to September precipitation averages 5–10 cm at most with peak
amounts along the crest of the southern California Peninsular Ranges, eastern San Bernardino Mountains, and the Sierra
Nevada divide from Mt. Whitney to Lake Tahoe. Amounts
decrease westward to 2 cm in the coastal plains, coast
ranges, and Central Valley because of stable air of Pacific
marine layer intrusions. Infrequent surges of very moist air
encourage outbreaks of thunderstorms throughout California
to the Pacific Coast. Decaying east Pacific tropical cyclones
enter California about once a decade, mostly south of Point
Conception. These storms are steered northward by southerly
upper air winds of the first seasonal troughs or cutoff lows
west of California, usually in September, and may produce
copious precipitation of 5–20 cm per day (Smith 1986).
In 1985, the Bureau of Land Management installed a
system of electromagnetic direction finders in the western
United States that record radiation emitted by cloud-to-ground
lightning strikes, with lightning located by triangulation
within 5 km resolution. The distribution of lightning strikes
reflects a combination of regional circulation and differences in air mass instability with terrain. Lightning detection
fluxes as high as 2–3 km2 yr1 occur in areas with greatest
summer precipitation, including the Peninsular Ranges,
eastern San Bernardino Mountains, the Sierra Nevada crest,
and high mountains to the east. Detections decrease to
0.5–1.0 km2 yr1 in desert basins and along the Pacific Coast
(GEOMET 1994; Minnich 2006).
Potential Evapotranspiration and Runoff
Potential evapotranspiration (PET) is the hypothetical moisture returned to the atmosphere from surface evaporation
and plant transpiration with unlimited water availability. It
is largely a function of temperature. PET varies locally along
gradients of wind speed and relative humidity. Annual rates
of 80–100 cm occur in areas with the coldest summers,
notably along the crest of the Sierra Nevada and the Pacific
Coast. Rates increase with mean annual temperature from
100 to 150 cm in the Central Valley and inland valleys of
southern California, and 150 to 250 cm in the southeastern
deserts (Kahrl et al. 1979). Because available water is finite,
actual evapotranspiration (ET) is much less than PET. Precipitation exceeds evapotranspiration during winter and
spring, resulting in runoff of 100 cm in the northwest
coast and northern Sierra, 40–100 cm in the Santa Cruz and
Santa Lucia Mountains, central Sierra Nevada, and southern
California Coastal Ranges. During summer PET exceeds ET
over the entire State. Summer thundershowers have limited
effect on soil moisture recharge. Lysimeter data for the
Sierra San Pedro Mártir in Baja California and the San Jacinto Mountains show that summer rain is countered by
high summer ET, with limited wetting of the root zone
(Franco-Vizcaino et al. 2002). Short-term rainfall intensities
are higher than in winter, increasing the runoff component
of the water budget.
In mediterranean climate, the available soil water is
dependent on recharge of the previous winter because plant
growth and transpiration demand are out of phase with seasonal precipitation distribution. In water surplus watersheds
(precipitation ET) soils are saturated virtually every rainy
season. Most precipitation variability is expressed in runoff,
with surpluses unavailable for ecosystem use. In waterdeficit watersheds (ET precipitation), the relationship
between precipitation and ET becomes stronger along a
decreasing precipitation gradient (Franco-Vizcaino et al.
2002). Hence, interannual variation in fire hazard is more
sensitive to precipitation in water deficit watersheds than
water surplus watersheds because the vegetation utilizes
most precipitation. Plant water supply in water deficit watersheds also depends on soil field capacities, which vary
greatly in mountainous terrain. Winter rains may saturate
steep slopes with low field capacities (e.g., 25 cm) virtually
every year, including subnormal rainy seasons. Vegetation
growing in basins on high field capacity soils may be insensitive to water deficits because of drainage from uplands.
Plant growth depends on seasonal temperature and precipitation cycles that vary with altitude and distance from
the Pacific coast. At lower elevations (1,000 m) mean winter temperatures range from 5C to 10C. Growth flushes of
mostly herbaceous ecosystems begin with the first winter
rains, usually in October in northern California and November/December in southern California; peak productivity
occurs in spring. Above 1,000 m growth flushes in chaparral,
woodlands, and conifer forests occur in late spring after
temperatures have warmed. In high elevation conifer forests,
growth occurs after snow melt, usually by early summer
(Royce and Barbour 2001a, 2001b). The onset of drought is
not significantly delayed compared to forests and woodlands
at lower elevations because moisture is rapidly depleted in
porous soils, especially in glaciated terrain. Snowmelt and
growth may be postponed to August in subalpine forests
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after wet winters. Summer thundershowers have limited
effect on soil moisture because PET exceeds precipitation by
an order of magnitude everywhere in the state (Franco-Vizcaino et al. 2002). The growing season along the immediate
Pacific coast also extends into summer due to cool temperatures and low transpiration demand.
precipitation (PON) reaches 150%–200%. Very strong El
Niños (e.g., 1982–1983, 1997–1998) produce above-normal
precipitation throughout the State. Precipitation during
weak El Niños shows a similar trend, but the PON is less
extreme than that for strong events. In La Niña/cold
episodes, precipitation tends to be above normal in northern California, normal from San Francisco to San Luis
Obispo, and below normal south of Point Conception.
Climatic Variability in the Instrumental Record
The mediterranean climate of California is well known for
interannual and interdecadal variability, especially total precipitation, because the region lies at the southern margin of
the jet stream. Recent investigations have correlated climate
variability in California with two coupled ocean-atmospheric
phenomena, the El Niño-Southern Oscillation and the
Pacific Decadal Oscillation (summarized in Minnich 2006).
E L N IÑO -SOUTH E R N OSCI LL ATION
The El Niño-Southern Oscillation (ENSO) comprises the
largest source of interannual variability in the troposphere
(Lau and Sheu 1991; Diaz and Markgraf 1992). Precipitation
variability in California is related to the interannual dislocation of the polar front jet stream along the Pacific coast by
ENSO. During ENSO events the weakening of trade winds
and eastward propagation of warm surface water masses
from the western Pacific Ocean (Kelvin waves) results in
above-normal sea surface temperatures (SSTs) in the central
and eastern equatorial Pacific Ocean. Theoretical studies
indicate that the release of latent heat and divergent outflow (poleward flux of angular momentum) from a region of
enhanced tropical convection in the central equatorial
Pacific Ocean result in dislocations of the Polar-front jet
stream, and anomalous rainfall patterns in North America
(Philander 1990; Lau and Sheu 1991).
During El Niño events, the atmosphere is characterized by
troughing of the jet stream in the Gulf of Alaska, blocking
(high pressure aloft) in western Canada, and an enhanced
southern branch jet stream across the southern United
States, with above-normal precipitation in California (Renwick and Wallace 1996; Hoerling et al. 1997). During La
Niña events, the jet stream lies poleward of California, with
weak onshore advection of moist surface westerlies into the
State. The recurrence interval of warm-phase El Niño events
is variable, but power spectra give a robust peak of 4 years.
La Niña events are equally infrequent, but SST “neutral”
years resemble La Niña events because the warmest equatorial SSTs still lay in the western Pacific.
Long-term statistical studies for California show that
annual precipitation correlates with ENSO (McGuirk 1982;
Ropelewski and Halpert 1986,1987; Schonher and Nicholson 1989; Cayan and Webb 1992; Montiverdi and Null
1997; Minnich et al. 2000). During El Niño events, precipitation departures increase equatorward along the Pacific
coast, with maximum values in southern California and
northern Baja California where the percentage of normal
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PACI F IC DECADAL OSCI LLATION
The Pacific Decadal Oscillation (PDO) is a rotating SST
anomaly around the North Pacific gyre that affects both precipitation and temperature in the western United States over
time scales of decades (Mantua 1999; Minobe 1999; Chao,
Ghil, and McWilliams 2000). Several models and mechanisms have been put forth to explain the oscillation, including small variations in solar irradiance influencing changes
in SST and zonal winds (Dettinger and White 2002), tropical
cloud cover (Wielicki et al. 2002), and variability in earthbound galactic cosmic rays (Mecurio 2003). Another mechanism is a rotating SST anomaly is the spin rate of the gyre
that occurs in two phases (Latif and Barnett 1994; Zhang,
Rothstein, and Busalacchi 1998; Barnett et al. 1999). In the
negative phase PDO, SSTs are above normal from Japan to
the Gulf of Alaska and below normal from California to New
Guinea. Reduced latitudinal temperature gradients result in
a northward shift of the jet stream, and a weakening of the
Aleutian low causing subnormal precepitation in California
(Cayan 1996; McCabe and Dettinger 1999). The negative
phase is associated with a “spin-up” of the north Pacific gyre
(increasing gyre rotation), with enhanced advection of warm
seawaters poleward off Japan. The dominance of high pressure over the entire north Pacific increases frictional wind
drag (wind stress curl), and the spin rate of the gyre, which
includes enhanced upwelling and colder than normal SSTs
along the California current. After ca. 10–25 years of the
spin-up mode, the increased gyre spin rate results in rotational advection of the warm pool to the North American
coast and then southwestward to the tropics (reduced
upwelling and warmer than normal SSTs along the California Current), with cold waters entering Japan and the Gulf
of Alaska (positive-phase PDO). These trends increase latitudinal temperature gradients across the Pacific.
The intensification and equatorward displacement of the
jet stream, as well as the associated enlargement of the Aleutian low, results in extratropical cyclones shifting equatorward
along the Pacific coast, causing above-normal precipitation
in California (Cayan 1996; McCabe and Dettinger 1999) and
reduced spin rate of the gyre. This mode also persists for one
or two decades when the “spin-down” phase is reversed by a
phase change to the negative mode. In another model, Hunt
(2001), arguing from a thermodynamic perspective, suggests
that the decadal oscillations are produced by El Niño cyclicity tied to global temperature change. Given earth–sun geometry, decades of global warming (e.g., 1910–1940, 1977–1999)
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are primarily expressed in increasing temperature in tropical
seas. Increased latitudinal temperature gradients generate
more frequent El Niños, which reduces warm water advection off Japan, further increasing latitudinal temperature gradients in a positive feedback. Decades of global cooling
(1940–1977) result in a reduced frequency of El Niños
(greater frequency of La Niñas) and enhanced advection of
warm water off Japan. Feedbacks produced by interannual
El Nino cyclicity may result in multidecadal oscillations at
high latitudes, perhaps varying the spin rate of the gyre
given in the gyre rotation model (Latif and Barnett 1994;
Gershunov and Barnett 1998). Cai, McPhaden, and Collier
(2004) show that the strength of El Niño cycles is related to
differences in the upwelling temperature of equatorial Pacific
waters, which shift on multidecadal cycles.
Minobe (1997, 1999) identifies four shifts (two cycles) in
the PDO, with each complete cycle interval spanning about
50–70 years. The last two shifts include a negative phase
from 1948 to1976 and a positive phase from 1976 to1999.
In California, the phase transition in the mid-1970s divides
a period of long-term drought the previous 25 years and
above normal precipitation since then, the latter period
having increased ENSO frequency. The two previous shifts
were a positive phase in 1925–1947 and a negative phase in
1890–1925. However, California precipitation records indicate that these phases were associated with less extreme wet
and dry episodes than those in the recent cycle. Over the
western United States, both negative phase shifts were associated with cold winters and springs, and both positive
phase shifts with warm winters and springs. Millar et al.
(2002) found acclerated tree growth in the Sierra Nevada
during warm phases of the PDO. Redmond et al. (1996)
show a positive correlation in growth of California blue oak
(Quercus douglasii) and precipitation. Evans et al. (2001),
using Sr:Ca ratios from coral in the equatorial Pacific correlated SSTs and the PDO since 1726. However, they found
that the strongest correlation is from the 20th century.
Analysis of instrumental records (1880–1994) and tree
ring data (back to 1700) shows a north–south sea-saw in
precipitation variability in western North America, pivoting
at 40N (Dettinger et al. 1998). Precipitation at California
latitudes are positively correlated with SOI which cause
southward displacements of precipitation distribution at
interannual and decadal time scales. The overall precipitation amount delivered to the west coast has little variation
both in the instrumental record and in tree ring records,
consistent with the century scale analysis of tree ring data
in the Sierra Nevada (Graumlich 1993).
Tertiary Climate Change and Development
of California Vegetation
Over geologic time scales, California has been both a biotic
repository and a hotbed of evolution and speciation. The
region has retained rich floral diversity in part because it
hosts temperate equable climates that were formerly exten-
sive across North America. Diversity is also linked with
remarkable environmental heterogeneity over a broad span
of latitude and over complex terrains. As stated by Axelrod
(1988), the taxa that contribute to the distinctive vegetation
of California have come from diverse sources and at different times during the Cenozoic; that is, the modern flora
comprises taxa that have divergent ancestral histories of
evolution, migration, and paleobiogeography (see Raven
and Axelrod 1978).
Knowledge of the paleobotany of the California flora is
vital to ecologists because the ancestral history of the
region’s biota, largely of Paleoarctic/Nearctic origin, permits
insight into such questions as adaptation, life traits, and
physiology; and most importantly it provides temporal context to modern distributions. The paleobotanical record in
California and adjoining regions is poor, as elsewhere in the
world, because the preservation of macrofossils is a rare
event in terrestrial environments. Plant microfossil (pollen)
records are invariably limited to wind-pollinated species,
and broadcast pollen dispersal compromises the spatial resolution of vegetation evidence.
Studies of paleofloras have invariably been coupled with
climate change. Until recently the biota was a primary
source of evidence in the reconstruction of past environments. However, the interpretation of climate and ecosystem change from paleobotanical evidence is confounded by
circularity because plants and climate are not treated independently. In addition, the most geographically widespread
species adapt to broad environmental diversity due to
genetic variability of populations and through physiological
plasticity, limiting precise inference of paleoclimate.
Besides, questions on plant migration and evolution should
be treated as legitimate processes on their own right (Sauer
1988; Davis and Shaw 2001).
The ability to treat biotic and climatic evidence independently has been bolstered by the recent emergence of
studies that infer environmental change from new data
sources and methods including stable isotope geochemistry,
global climatic models (Earth surface/atmospheric circulation models), Milankovich astronomical cycles, and traditional geomorphic evidences. Recent summaries on the use
and limitations of proxy data in climate reconstructions are
given in Lowe and Walker (1998), Bradley (1999), Graham
(1999), and Ruddiman (2001).
This section reviews global climate change, local plate
tectonic history, and the development of the California
flora since the late Cretaceous. Paleobiogeographic and climatic evidences are then compared to address how climate
has influenced the paleobiota leading to modern vegetation, primarily with the intent to develop questions that
may be addressed in future research.
Cenozoic Global Cooling
Since the late Cretaceous (100 ma), the Earth has shifted
from a “greenhouse” climate to an “icehouse” climate. The
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most compelling evidence is the global deep sea record of
paleotemperatures obtained from 18O/16O isotope ratios
within the mineral walls of benthic marine invertebrates in
which more 18O is taken up in proportion to 16O as water
cools (Savin 1977). Oxygen isotope ratios are also directly
influenced by the growth and decay of ice sheets which
preferentially take up 16O, making higher 18O/16O seawater
ratios (16O evaporates more easily into the atmosphere than
18
O). With respect to carbon, 12C is preferentially taken up
in photosynthesis, so that organic carbon is enriched in 12C
relative to 13C. When sea level is high during interglacial
periods, the storage of organic matter in terrestrial biomass
and in sediments of flooded continental shelves result in
heavy or positive 13C ratios recorded by marine organisms.
During glacial periods, when low sea levels expose continental shelves, organic matter enriched in 12C is released
into the oceans, producing light (negative) 13C ratios. However, the relationship between climate and benthic 13C/12C
ratios is less clear than for 18O/16O ratios because deep-sea
isotopic records often shows a positive shift in 13C during
periods of ice sheet growth, perhaps reflecting enhanced
ventilation of deep seawater masses.
Sea floor cores have shown that deep-sea temperatures
reconstructed from oxygen isotope ratios have cooled over
most of the Cenozoic (summaries in Abreu and Haddad 1998;
Zachos et al. 2001). Oceanic temperatures first increased from
the Cretaceous-Tertiary boundary to the Eocene climatic optimum, with benthic temperatures exceeding 15C that was
associated with a global carbon isotope excursion (CIE) dated
at 55.5 Ma. Planktonic ocean temperatures at high latitudes
reached 20C. Temperatures remained warm into the early
Eocene. The CIE suggests that the ocean sustained a rapid and
large input of isotopically light carbon due possibly to massive
release of biogenic methane from methane hydrates in continental shelf sediments (Rea 1998; Bains, Corfield, and Norris
1999; Katz, et al. 1999). Methane may have come from deeper
ocean sources (Thomas et al. 2002). Benthic ocean temperatures have since experienced stepwise decreases to 10C in the
late Oligocene, 6C–7C in the middle Miocene, and modern
levels of 2C–4C in the Plio-Pleistocene. The 13C/12C ratios
show similar excursions in the Cenozoic, as 13C ratios are positive from the Paleocene to Miocene and negative in icehouse
climates in the Pliocene and Pleistocene.
The changing global climate is related to plate tectonics,
which influence the latitudinal distribution of continents
and oceans, changes in global sea level, and orogenic production of mountain ranges and plateaus, all of which affect
the planetary absorption and redistribution of solar energy,
as mediated by shifting ocean and atmospheric circulation
(summary in Ruddiman 2001). Changing latitudinal temperature gradients also affect cloud cover, precipitation distribution, and the abundance of the greenhouse gases H2O
and CO2 and the planetary radiation budget. Milankovitch
orbital cycles, that is, the precession of the axis (23 ka), the
Earth’s tilt (41 ka), and ellipticity (100 ka) have affected
global climate by causing latitudinal and hemispheric redis-
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tribution of insolation, with greatest impact during Quaternary glacial cycles (Zachos et al. 2001; summarized in
Bradley 1999; Ruddiman 2001).
During the Cretaceous and early Tertiary, the warmth of
the deep ocean resulted in warm polar climates where there
was free mixing of surface and warm deep waters (reduced
thermal stratification with depth). Equatorial sea surface
temperatures were apparently similar to those today because
the potential for increasing SSTs in “greenhouse” climates is
countered by high evaporation rates. Since tropical SSTs
were comparable to that today (Zachos et al. 2002), latitudinal ocean and atmospheric temperature gradients must have
been weaker (15C) than at present (30C). The increase in
atmospheric moisture content in the subtropics and midlatitudes also increased the poleward latent heat flux, thereby
reducing the strength of the jet stream and eddy (cyclone)
activity (Pierrehumbert 2002). Sea-level maxima due to rapid
sea floor spreading rates resulted in shallow seas covering
extensive continental areas, including a midcontinental seaway from the Gulf of Mexico to the proto-Arctic Ocean. The
reduction of continental surface exposure and deserts may
have decreased planetary albedo. In the southern hemisphere,
the north–south configuration of the Antarctic–Australian
continent encouraged poleward heat transport both in the
ocean and atmosphere at rates sufficient to support forest
ecosystems and a marsupial mammal fauna in Antarctica
(e.g., Woodburne and Zinsmeister 1982; Askin 1992).
Geomorphic evidences suggest that the planet was free of
ice sheets.
Global cooling began in the Eocene as new plate tectonic
arrangements reduced meridional heat transport to polar
regions. The northward rifting of Australia left Antarctica
centered over the South Pole, the resulting zonal ocean and
atmospheric circulations leading to rapid cooling and onset
of glaciation in Antarctica by the Oligocene. This contributed
to a rapid decline in global deep-sea temperatures due to
onset of cold bottom waters characteristic of modern oceans.
By the early Miocene, global cooling and intensification of
the poleward heat gradient both reduced the poleward latent
heat flux and increased the strength of the westerlies and
eddy (cyclonic) activity (Pierrehumbert 2002). Northern
hemisphere continental climates cooled with the uplift of
the Rocky Mountains (Laramide orogeny) and the Tibetan
plateau. These high terrains served to isolate Northern
Hemisphere landmasses from maritime airflows. By the
Miocene, the opening of the Drake Passage between South
America and the Antarctic Palmer Peninsula intensified
strong west–east ocean-atmospheric circulations surrounding
Antarctica, with permanent continental-scale glaciation beginning ca. 10 ma. The progressive cooling of deep-sea temperatures led to upwelling of ever-cooler waters off the west
coasts of continents, including California. The resulting
stabilization of eastern Pacific air masses reduced summer
precipitation over western North America. The strengthening
westerlies and cyclonic activity with global cooling increased
winter precipitation, especially after the mid-Miocene.
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The closing of the Panama Isthmus led to the establishment of the Atlantic Gulf Stream and high latitude warm
ocean source of heat and water vapor for development of
Northern Hemisphere ice sheets beginning ca. 2–3 ma. The
Arctic ice pack also developed about this time. Glacial cycles
are broadly phased with Milankovitch orbital cycles, with
the buildup of ice sheets during northern hemisphere solar
minima in summer and interglacials during solar maxima in
summer. Benthic 18O/16O ratios show that early glaciations
were phased with 23 ka precession and 40 ka axial tilt cycles
associated with precession and the axial tilt until ca. 850 ka
after which glacial cycles became dominated by 100 ka
orbital ellipticity cycles. Ice ores of the Greenland and
Antarctic ice sheets reveal that the abundance of the greenhouse gases CO2 and CH4 parallel 18O/16O ratios, with lower
abundance during glacial cycles. The ice sheets, Arctic ice
pack, and extensive seasonal snow cover increased global
surface albedo. Rapid erosion rates due to the uplift of the
Tibetan Plateau and other young mountain ranges may
have contributed increased chemical weathering and lowering of atmospheric CO2.
Many aspects of Quaternary climate change are not
phased with Milankovitch cycles (Lambeck, Esat, and Potter 2002; summary in Bradley 1999). The growth of ice
sheets couples with warm Atlantic Ocean temperatures
and increased water vapor transport onto the North America and Eurasia at scales of millennia. North Atlantic SSTs
fluctuate with phase shifts of the oceanic conveyer belt
that responds to ice sheet dynamics. Within a glacial cycle
are Dansgaard-Oeschger and Heinrich events produce
alternating episodes of severe and quiescent weather at
millenial scales. They appear to be phased with ice sheet
flow systems that affect the production of icebergs and the
intensity of the Gulf Stream heat flux in the North
Atlantic.
In the last glacial cycle, the North American ice sheets
first developed near the Arctic Ocean including the Canadian Archipelago at ca. 100 ka. Ice sheets built southward
to cover most of Canada (Laurentide ice sheet) during the
Milankovitch solar minimum at 75 ka. Ice sheet volume
fluctuated thereafter, reaching maximum volume in the
late glacial maximum at 30–18 ka. Marine oxygen isotope records show that the northern hemisphere ice
sheets thinned and retreated between 16.0 and 11.0 ka, in
phase with a Milankovitch solar maximum at 12.0 ka.
This trend was reversed briefly in the Dryas advance at
12.5 ka, which was followed by the final collapse of Laurentide ice sheet by 8 ka. The modern Holocene climate
began at ca. 11 ka, although it can also be viewed as one
of many brief anomalous interglacials inasmuch as global
climate has been in a glacial mode ca. 90% of the time
since 850 ka.
The climate of California during the Tertiary shifted from
tropical to mediterranean as global climate shifted from
“greenhouse” to “icehouse” states. Because of small latitudinal temperature gradients required for poleward heat flux,
the influence of the jet stream during the late Cretaceous
and Paleogene was weak, compared to the present. Warm
oceans along the Pacific coast were a source of moist unstable air masses and year-round convective precipitation,
either from winter frontal systems or moist trade wind circulations in summer. Because high accretionary terrains of
Mexico had assembled by the late Cretaceous (OrtegaGutierrez, Sedlock, and Speed 1994), California was apparently shadowed from Atlantic tropical air masses over most
of the Cenozoic. Low seasonal temperature cycles and
equable climates, and floras with tropical forest ancestry
imply frost-free winters. Global cooling and eventual continental glaciation in Antarctica by the Miocene, and the
northern hemisphere polar regions by the Plio-Pleistocene,
resulted in ever-cooler climates in California and western
North America. The cooling of deep seas and upwelling of
ever-colder water along the Pacific coast resulted in the
development of summer drought by the Miocene. California’s climate has fluctuated with glacial–interglacial cycles
since the onset of the Pleistocene, precipitation greater than
present during glacials and climates comparable to that at
present during interglacials.
Plate Tectonics and Topography
of the Western United States
The paleobiogeography of floras in the western United
States has been influenced by ever-changing topography
along the western margin of the North American plate
(summary in Dickinson 1981). A passive margin in the
early Paleozoic, similar to present plate tectonic arrangements in the eastern United States, had California under
sea level. The region entered a Japanese-style subduction
complex during the remainder of the Paleozoic and into
the late Triassic, and an Andean-style subduction phase
from this time until the Oligocene-Miocene time boundary. Since the late Cretaceous, orogenies have led to the
development of the Sierra Nevada, south coast batholiths,
and the accretionary terrains of central and northern California Coastal Ranges, including the Franciscan sedimentary complex. The Andean-phase subduction was
related to the Laramide orogeny and uplift of the Rocky
Mountains that ended in the Eocene. By the early
Miocene, volcanic arc subduction was terminated by the
encounter of western North America with the eastern
Pacific spreading ridge, which led to the development of
the San Andreas transform system. During this period,
basin and range extensional terrains developed in the
Great Basin. Neogene tectonics (related to the San
Andreas transform) throughout the coast of California
contributed to the rapid rise of the Sierra Nevada and
South Coast Ranges beginning in the late Miocene and
Pliocene. The uplift of the Pacific Coast Ranges resulted
in a new topographic axis, with intervening elevated surfaces making up the basin and range province and the
Colorado Plateau.
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Paleovegetation
The following is a brief review of the paleobotanical record as
summarized in Graham (1999) Late Cretaceous and Cenozoic
history of North American Vegetation, and Axelrod (1988) in the
previous edition of Terrestrial Vegetation of California which
emphasizes paleofloras in the western United States and California. These sources should be consulted for a more complete
account of the evidence than presented here, as well as the
primary literature. The California record is discussed in the
broader context of North Amercian paleobiogeography
because the California biota is an outcome of the evolution,
immigrations, and extinctions operating at continental scales.
TA B L E 2.2
Classification of Tertiary Forests
Tropical Forest
A broad-leaved evergreen single-tiered open canopy
vegetation, leaves mostly entire-margined, thick
textured, few drip tips. Climate subhumid, little
seasonality.
Tropical Rain Forest
Broad-leaved, evergreen, multistratal; drip tips, lianas
and buttressing; leaves sclerophyllous, mostly
mesosphylls, dominantly entire-margined leaves.
Climate humid, little seasonality.
LATE CR ETACEOUS TO EAR LY EO CE N E
During the late Cretaceous and early Paleogene (Paleocene
and Eocene), virtually the entire North American continent
was covered by forest. The floristic composition of the vegetation is not well known. Many Paleocene through middleEocene angiosperms assigned to modern taxa in older literature are now thought to represent extinct or intermediate
forms (Graham 1999). A recent approach has been to
describe the vegetation using the method of Wolfe (1971),
using leaf size, presence of drip tips, percentage of species
with entire leaf margins, and whether species were evergreen
or deciduous. A classification of broad vegetation formations,
using this method, is followed in this review (Table 2.2).
In the late Cretaceous, tropical forests extended from the
southeastern United States toward Greenland. These forests
were a mixture of tropical genera (i.e., lineages allied to those
in modern tropical floras), and early representatives of Taxodiaceae similar to Sequoia and Metasequoia, as well as Cupressaceae and palms. Paratropical forests covering the western
United States were replaced by notophyllous broadleaved
evergreen forest to the north. Notophyllous forests contained
Equisetum, Cycadophytes, and Ginkgophyte representatives
from earlier Cretaceous vegetation, as well as Araucariaceae
and evergreen Taxodiaceae. Broadleaved deciduous forest
occupied the high latitudes. The origin of deciduousness in
angiosperms, which extends back to the mid-Cretaceous, is
unclear. Axelrod (1991) hypothesized that deciduousness in
angiosperms began when they encountered seasonal dry climates as they migrated from their ancestral home in Gondwana in the late Mesozoic. Deciduousness preadapted plants
to polar nights when these taxa reached high latitudes of
Eurasia and North America. Alternatively, warm Arctic climates, combined with polar day–night cycles, may also have
served as a “birth place” of novelties such as deciduousness
that led to the evolution major plant groups that later radiated to lower latitudes (Hickey et al. 1983).
In the Paleocene, tropical rain forests existed in the
southeastern United States, with deciduous forest in the
Appalachian uplands. Notophyllous broad-leaved evergreen
forest, which occurred along coast of the North Atlantic land
bridge with Europe gave place to polar broadleaf deciduous
forest toward the pole. A series of floras in marine sediments
54
Paratropical Rain Forest
Floristically allied to tropical rain forest,
predominantly broad-leaved evergreen with some
deciduous, woody lianas diverse and abundant,
buttressing present; mostly entire-margined leaves.
Precipitation seasonal, but no extended dry season,
some frost.
Subtropical Forest
Sclerophylls abundant, few lianas, no buttressing,
predominantly a broad-leaved evergreen forest
with some conifers and broad-leaved deciduous
elements; about half species with entire-margined
leaves. Frost present but not severe, seasonal rainfall,
cool winters.
Notophyllous Broad-leaved Evergreen Forest
Some broad-leaved deciduous trees, conifers not
common, abundant woody climbers, buttressing absent,
sclerophyllous, no drip tips, small mesophyll
(notophyll) leaves. Cool winters, warm summers.
Warm Temperate Forest
Broad-leaved deciduous forest, some conifers, and
broad-leaved evergreens, but not dominant. Minority of
species with entire-margined leaves. Winters below
freezing, warm summers.
Polar Broad-leaved Deciduous Forest
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An extinct mesotherm to microthermal moist
forest type related to the modern deciduous
forest formation. Leaves large, thin-textured,
and seldom entire. Winters below freezing, mild
summers.
SOURCE :
Graham (1999, 149), after Wolfe (1971).
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of the midcontinental seaway record diverse ferns, as well as
Ginkgo, Taxodiaceae (Glyptostrobus, Metasequoia), araucarias,
palms, and numerous notophyllous angiosperm leaves in
Juglandaceae, Betulaceae, Cornaceae, Fagaceae, platanoids,
Ulmaceae, Moraceae, and Lauraceae. In the Northwest Territory of Canada, fossil floras record Cupressaceae, Taxodiaceae, Picea, Pinus, Acer, Alnus, Betula, Ericaceae, Juglandaceae, Liquidambar, and lineages similar to Quercus, Tilia,
and Ulmus.
Early Eocene floras record paratropical forest extending
across the Great Plains to the eastern slope of Rocky Mountains. In the western United States, coastal areas were occupied by tropical to paratropical rain forest, which gave place
to notophyllous broad-leaf evergreen forest and then to
polar broadleaved deciduous forest north toward Alaska.
The central Rocky Mountains contained mesothermal sclerophyllous broad-leaved evergreen tropical to paratropical
rain forest and mixed mesophytic elements with Asian
affinities, including some taxa now present in tropical
northern Latin America. These forests had a few temperate
elements including Acer, Alnus, and Populus. There are few
fossil floras in California and adjoining regions from the
late Cretaceous to the early Eocene. Based on floras farther
north along the Pacific coast, Axelrod (1988) hypothesized
that the region was covered with temperate rain forest grading southward into subtropical forest. Taxa recorded in Oregon and Washington include tree ferns, palms, cycads,
large-leaved evergreen trees of tropical and subtropical families, particularly dominated by members of Anacardiaceae,
Burseraceae, and Caesaldaceae.
M I DDLE EO CE N E AN D OLIGO CE N E
By the middle Eocene, the southeastern United States was
covered with semideciduous tropical dry forest. Fossil floras
near the Appalachians indicate the forest on lower slopes
consisted of warm-temperate deciduous elements of Fagaceae
and Juglandaceae with more tropical elements in lower
Appalachians and coniferous elements (Picea, Tsuga) in the
high Appalachians. Along the Gulf coastal plain, tropical dry
forest was replaced by warm temperate deciduous forest, with
conifer forest persisting in the highlands. Important genera
include Fagus, Acer, Quercus, Castanea, hickory, and Taxodium,
with many Asian exotics. This formation extended as far
north as Ellesmere Island (lat. 81N), where there was also
Larix, Picea, Tsuga, and Betula. Semideciduous tropical dry
forest extended from the Gulf Coast to the southern Rockies
and southern California. Forested areas in the Great Plains
gave way to less woody biomass and increasing abundance of
grasses. The Rocky Mountain highlands from British Columbia to New Mexico had conifer forests that were apparently
ancestral to modern western montane conifer forests (Picea,
Pinus, Larix, Thuja, Tsuga, Metasequoia, Ginkgo, Sequoia, and
Abies). These forests had a deciduous component consisting
of Alnus, Carya, Juglans, Betula, Castanea, Quercus, Populus,
and Acer. There are no records in coastal California from this
period, but the vegetation is believed to have been a tropical
rainforest because similar vegetation was present farther
north along the coast. Floras from this period have the last
appearances of several tropical forest elements, for example,
Milfordia (Restionaceae; Graham 1999), which were replaced
by semiarid tropical groups in Bursera, Pachysander-Sarcococca,
Restionaceae, Fremontodendron, and Juglandaceae. Eocene fossil records suggest an early development of a chaparral,
woodland, and savanna component to the vegetation in the
Southwest. The Susanville flora has Lauraceae, Magnolia,
Ulmoideae, and notophyllous-to-megaphyllous broad-leaved
evergreens. The Green River flora has Ephedra, Ocotea, Rhus,
Quercus, Mahonia, and Rhamnus. It was hypothesized by Axelrod (1988) that shrubland, chaparral, woodland, and savanna
in the Neogene were derived from these Eocene evergreen
sclerophyll components.
There are few fossil records in North America from the
Oligocene. In the eastern United States, the most important
fossil flora is the Brandon lignite in Vermont which contained a mixture of conifers, deciduous hardwoods, and
evergreen hardwoods in Pinus, Glyptostrobus, Pinus, Rhus,
Ilex, Rhododendron, Vaccinium, Castanea, Fagus, Quercus, Liquidambar, Juglans, Carya, a Persea-type, an Ocotea-type, Magnolia, Morus, Rhamnus, Tilia, Ulmus, and Vitus. At Devon
Island Canada (75N), the vegetation consisted of a mixed
conifer deciduous forest with Larix, Picea, Pinus, Tsuga,
Alnus, Betula, Ulmus, Coryles, Ericales, Glyptostrobus, Metasequoia, Abies, and Thuja. The Great Plains had become a
savanna by the Oligocene, especially in rain shadows east of
the Rocky Mountains. In the west, the Creede flora in
southwestern Colorado records a mixed hardwood-conifer
forest, mountain mahogany chaparral, and perhaps pinyonjuniper woodland. Fossil floras in Oregon contain temperate broadleaved deciduous forest. The coastal strip contained subtropical elements and broad-leaved evergreens
(Annona, Nectandra, Ocotea). Northwestern Canada was covered with conifer forest of Picea, Pinus, Tsuga, Alnus, Betula,
Sequoia, and Quercus.
M IO CE N E TO EAR LY P LE I STO CE N E
By the early Miocene, tropical dry forest and much of the
notophyllous broad-leaved evergreen vegetation covering
the eastern United States had disappeared, leaving deciduous forest and pine woods. According to Axelrod (1988),
conifer forests in the Rocky cordillera are widely recorded in
Nevada, Idaho, and the Rocky Mountains, with Chamaecyparis, Picea, Pseudotsuga, and Tsuga. In the interior Pacific
Northwest, these genera were associated with Metasequoia,
Sequoia, and Thuja, and in the southern Great Basin, with
Pinus and Abies. The Rocky Mountain forests still had a rich
flora of deciduous hardwoods (Betula, Carya, Fagus, Ostrya,
Quercus, Sassafras, Tilia, Ulmus, Liquidambar, Acer, Fraxinus,
Populus, and Alnus). Many hardwoods have descendant
species now occurring in western United States, but a few
occur in the eastern United States and eastern Asia. Many
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Miocene conifers in the northern Rocky Mountains and
Idaho are allied to the modern north coast forest, whereas
Miocene floras east of the Sierra Nevada were composed of
species whose nearest descendants contribute to modern
Sierran mixed-conifer and subalpine forests.
Miocene floras in Oregon, Washington, west Idaho,
Nevada, the central Great Basin, near Coalinga CA, and in
the western Sierra Nevada (late Miocene) contained ancestral elements of modern mixed evergreen forest (what Axelrod [1988] called oak-laurel forest). Evergreen trees included
Arbutus, Chrysolepis, Lithocarpus, Quercus, Magnolia, Persea,
and Umbellularia. These floras also contained deciduous
hardwoods still extant in California (Acer, Alnus, Quercus
(kelloggii), Prunus, Rhamnus), as well as evergreens (Cedrela,
Ilex, Persea, Oreopanax), and hardwoods similar to taxa now
in the eastern United States or eastern Asia (Ailanthus, Carpinus, Castanea, Carya, Fagus, Liquidambar, and Sassafras). It
was assumed that the Sierra Nevada had only attained a low
elevation in the Miocene and that the hills contained a rich
deciduous forest with numerous evergreen dicots, an assemblage that apparently extended northward to the western
Cascades. Several fossil localities in central and southern
California (e.g., Tehachapi, Mint Canyon) record an oaklaurel forest dominated by Quercus species ancestral to modern California species and exotic species now in Mexico and
the eastern United States, as well as a diverse assemblage of
laurels in Nectandra, Ocotea, Persea, and Umbellularia (Axelrod 1988). These floras also record Arbutus, Ficus, Ilex,
Lyonothamnus, and evergreen sclerophyllous shrubs in Cercocarpus, Prunus, Quercus, Rhus, Arctostaphylos, Ceanothus,
Mahonia, Rhamnus, Fremontodendron, Heteromeles, and Malosma. The vegetation also contained ancestral species to
modern pinyon-juniper woodland, cypress, and tropical
components in Acacia, Bursera, Ficus, and palms. A middleMiocene mountain flora in the Santa Monica Mountains
records a mixture of conifers, hardwoods, and chaparral
including Pinus (ponderosa), Calocedrus, Pseudotsuga, Quercus,
Cercocarpus, and Rhamnaceae (Stadum and Weigand 1999).
Floras from the Middle Miocene through the Pliocene in
the eastern United States reveal a modernization and expansion of mixed mesophytic deciduous forest in association
with the disappearance of Asian and Neotropical elements.
An increasing number of conifers mixed with deciduous
forests in the north. A boreal forest developing in Canada
still had Asian exotics and deciduous hardwoods (Glyptostrobus, Ptercarya, Tilia, Ulmus), but these taxa disappeared
by the Pliocene. Grassland expanded at the expense of
deciduous forest in central part of country, but along the
eastern slopes of the central and southern Rocky Mountains
grassland gave way to pinyon-juniper woodland.
In the Pliocene, conifers descendent from Eocene
Cordilleran forests reached the coast of Washington, Oregon,
and northern California. A rich coast conifer forest is
recorded near Santa Rosa in the late Pliocene with Abies,
Picea, Tsuga, and Rhododendron. The late Pliocene Sonoma
flora contained Pinus (ponderosa), Pseudotsuga, Quercus spp.
56
and Q. wislizenoides, and chaparral of Arctostaphylos, Cercocarpus, and Heteromeles. Along the coast nearby, a Sequoia forest
grew with Alnus rubroides, Chrysolepis, Ceanothus, Lithocarpus,
Mahonia, and Umbellularia. A Plio-Pleistocene flora of mixed
conifer forest with broad-leaved sclerophyll vegetation was
recorded near Santa Clara. Late Miocene floras in western
Nevada record Abies bractiata with Acer, Arbutus, Lithocarpus,
Pinus ponderosa, Quercus chrysolepis, and Q. wislizenii. Early
Miocene mixed conifer forest and mixed evergreen forests
disappeared from the lowlands of western Nevada by early
Pliocene, replaced by oak-juniper savanna and pinyonjuniper woodland. Fossil floras contain Quercus (arizonica,
wislizenii) and chaparral (Cercocarpus, Rhamnus, Rhus, Arbutus
(arizonica), and Rhus). By late Pliocene, the chaparral component migrated southward to southern California and Arizona, apparently leaving behind pinyon-juniper woodlands
with an understory of subshrubs. An oak woodland savanna
occurred on the west flank of the Sierra Nevada by the end of
Miocene (Remington Hill flora) with Quercus, Arctostaphylos,
Ceanothus, and Prunus. The first oak woodlands in coastal
California were recorded in the early Pliocene.
There is evidence that mixed-evergreen and mixed-conifer
forests were present in the Sierra Nevada during the Pliocene.
Floras in the central part of the range record Quercus (Q.
chrysolepis, Q. agrifolia, Q. engelmannii), Lithocarpus and a few
remains of Abies and Sequoiadendron. Evergreen sclerophyllous vegetation was present in the San Francisco Bay in
the middle Pliocene (Arbutus, Heteromeles, Lithocarpus,
Lyonothamnus, Quercus, Umbellularia), associated with deciduous elements (Acer, Alnus, Amelanchier) and chaparral
species of Arctostaphylos and Rhamnus. The rich Mt. Eden
flora near Beaumont (5.3 ma; Axelrod 1937, 1950) contains
taxa representing a large range of environments due apparently to transport of organic material. The flora has members
of desert scrub and woodland (Baccharis, Cercidium, Chilopsis,
Ephedra, Ficus), chaparral (Arctostaphylos, Ceanothus, Cercocarpus, Prunus, Heteromeles, Quercus, Rhamnus, Rhus, Pinus (attenuata), Cupressus (forbesii), Malosma), woodland (Pinus, Arbutus, Juglans, Persea, Platanus, Populus, Quercus), and montane
conifer forest (Pinus, Pseudotsuga premacrocarpa, Quercus). The
Anaverde and Ricardo floras in southern California record
oaks and palms. The closed-cone conifers Pinus radiata and P.
muricata were present in middle and early Pliocene in interior
California, but their ancestral history is otherwise unknown.
Other members of Pinus sect. oocarpae now live in central and
southern Mexico. The pines in sect. oocarpae are believed to
have evolved in Mexico/Latin America in the Miocene (Millar 1999). A limited fossil record suggests that the California
members of this section have become increasingly restricted
to disjunct portions along the Pacific coast (Axelrod 1980,
1981). P. sabiniana is found in the southern California floras
in the middle Pliocene Mt. Eden flora, the late Pliocene Pico
flora, and a Plio-Pleistocene flora near Santa Paula.
There are few Tertiary records of desert scrub or coastal
sage scrub because soft leaves are unlikely to contribute to
the fossil record. Axelrod (1988) proposed that desert scrub
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may have been derived from two sources: east Asian desert
elements (Artemisia, Atriplex, Ephedra, Kochia, Suaeda) that
reached California via the Bering Strait in the Pliocene
before full glacial climates were established, and softstemmed shrubs with a sage habit (Chrysopsis, Chrysothamnus, Ericameria, Eriogonum, Salvia, Tetradymia, Salazaria) that
may have originated in the southwestern deserts. Desert
microphyll woodlands of the Sonoran Desert (Acacia, Bursera, Euphorbia, Pachycormus) are tropical lineages apparently
derived from Mexico.
During the Pleistocene, deciduous and conifer floras in
the eastern United States experienced multiple extirpations
and interglacial poleward invasions. Tertiary conifer ancesters in the Appalachians spread to lower elevations and
integrated with deciduous hardwoods in the northeastern
U.S. Boreal forest conifers and hardwoods that grew mostly
in northern Rocky Mountains during the Tertiary (Abies,
Chamaecyparis, Picea, Pinus, Pseudolarix, Thuja, Tsuga, Betula,
Populus) expanded into the Canadian lowlands. Extensive
palynological records show that during the late glacial maximum, boreal forest extended southward into eastern
United States in advance of ice sheets. Temperate hardwoods extended southward to the Gulf Coast and the Neovolcanic axis near Mexico City where deciduous forest genera still occur. During the Holocene boreal forests retreated
to the Canadian border and were replaced by modern deciduous forest. Grasslands persisted in the Great Plains during
both glacials and interglacials.
There are few macrofossil floras in California from the
early Pleistocene. Most noteworthy is the Soboba flora near
San Jacinto (Axelrod 1966) which contains 41 species from
several plant communities, including riparian woodland
(Acer, Cornus, Fraxinus, Platanus, Populus, Quercus), bigcone
Douglas fir-coulter pine forest (Pseudotsuga macrocarpa,
Pinus coulteri, Acer, Amorpha, Arbutus, Quercus chrysolepis,
and Q. wislizenii), mixed-conifer forest (Pinus ponderosa,
Pinus lambertiana, Calocedrus decurrens), and chaparral
(Ceanothus, Cercocarpus, Garrya, Mahonia, Prunus, Quercus,
Rhus, and Cupressus forbesii).
Pollen assemblages from Lake Tahoe with a minimum date
of 1.9 ma record a mixed-conifer forest pollen rain with
greater abundance of undifferentiated Taxaceae–Cupressaceae–Taxodiaceae (Calocedrus decurrens?) and lower Abies
and Quercus than present (Adam 1973). A core from Owens
Lake that extends back to 800 ka (Litwin et al. 1995) shows
Juniperus increasing relative to Pinus during glacials and conversely during interglacials. The increase in Juniperus may
reflect the establishment of the genus near the site, as in the
late glacial maximum. The peak abundances of Pinus may be
due to the decline of Juniperus in a background of continuous
pollen flux of white pine species including P. monophylla,
P. lambertiana, P. monticola, P. flexilis, and P. balfouriana, all of
which may have expanded their ranges in the Sierra Nevada
and White-Inyo Mountains during interglacials (Woolfenden
1996). Sediment cores from Clear Lake in the North Coast
Range show that pollen was dominated by Pinus in late
glacial and Quercus during the Holocene, with a large shift
toward Quercus at ca. 15 ka (Adam et al. 1981; Anderson and
West 1983). Today, the region is dominated by an oak woodland (Quercus lobata, Q. douglasii, Q. wislizenii) and Pinus
sabiniana. It is unclear whether the dominance of Pinus
pollen in full glacial climates represents Pinus sabiniana
foothill woodland or P. ponderosa mixed-conifer forests.
Late Pleistocene and Holocene
Most late Quaternary vegetation evidence in California comes
after the last glacial maximum (25 ka). Climatic and vegetation changes are unavoidably discussed together at this point
because the Quaternary has been defined by climatic events.
Emphasis here is given to California studies as summarized
from pollen and packrat midden studies by Woolfenden
(1996), Spaulding (1990), and Van Devender (1990).
LATE P LE I STO C E N E
GCMs for the late glacial maximum (LGM; 21–16 ka,
Bartlein et al. 1998) show the equatorward displacement of
the jet stream and moist surface westerlies to lat. 35N along
the Pacific Coast by the Laurentide ice sheet. Sea surface
temperature variations in the western and eastern Pacific,
using magnesium/calcium ratios in foraminifera suggest
that a persistent El Niño-like pattern existed in the tropical
Pacific during full glacial climates (super ENSO; Koutavas
et al. 2002; Stott et al. 2002). A super ENSO contributed to
the southward displacement of the jet stream and precipitation along the Pacific coast. High-surface albedos of the Laurentide ice sheet maintained a strong jet stream in summer
over the western United States.
Evidences for cooler wetter climates in California include
glaciation along the Sierra Nevada’s crest and pluvial lakes
in the Great Basin. Glacier equilibrium line altitudes (ELAs)
were 1,000 m below the ELAs of modern cirque glaciers in
the Sierra Nevada (Broecker and Denton 1990; Burbank
1991; Dawson 1992, 68). The lowest glacial cirques on
windward mesic mountain escarpments range from 2,000 m
in the Siskiyou-Cascades, to 2,400 m in the northern Sierra,
to 2,700 m at Yosemite, and to 3,000 m in the San
Bernardino Mountains of southern California. Cosmogenic
dating of terminal moraines indicate that mountain glaciers
in the San Bernardino Mountains reached maximum extent
20.0–16.0 ka, recessional moraines correlating with the
Dryas at 12.0 ka (Owen et al. 2003), in phase with glacial
advances in Yellowstone National Park (Licciardi et al. 2001)
and the Sierra Nevada (Benson et al. 1998). Carbonate-free
clay-size fraction of Owens Lake sediments indicate that
the crest of the Sierra Nevada was continuously glaciated
since at least 52.0 ka, the greatest extent occurring during
the Tioga advance, 24.5–13.6 ka (Benson et al. 1998). Pluvial
lakes in the Great Basin were supported by reduced evaporation rates due to cool summers that were associated with
a strong summer jet stream (Spaulding 1990; Bartlein et al.
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1998). Pluvial lakes increased after 32.0 ka and reached
maximum extent between 15.0 and 13.5 ka (Benson et al.
1990), then dried or decreased to modern levels by 10 ka
(Street and Grove 1979). The California deserts remained
semiarid to arid due to their location in rain shadows. Drill
cores at five sites along the floor of Death Valley show only
periodic lacustrine deposits, indicating that Lake Manley
(fed largely by the Armagosa River) was a small shallow lake
over most of the Pleistocene (Anderson 1998).
Pollen and Packrat midden studies in the Sierra Nevada
reveal profound vegetation shifts during the LGM (summary in Woolfenden 1996; Anderson 1990, 1996). Late glacial pollen sites at 1,500–2200 m in the Sierra Nevada
(Nichols Meadow, Nelder giant sequoia grove, Balsam
Meadow, Exchequer Meadow) record mostly Artemisia and
Poaceae, which is interpreted to reflect a sagebrush steppe
vegetation. Whether modern alpine herbs contributed to
this assemblage or whether the landscape was barren
with exotic Artemisia/Poaceae pollen passing through is
unknown. Other genera recorded include limited amounts
of Pinus and Cupressaceae. These sites, which lie only 500 m
below full-glacial ELAs of the Sierra Nevada ice sheet, were
apparently near or above the treeline. (Modern treelines of
3,000–3,500 m lie 500 m below modern glaciers.) Packrat
(Neotoma) middens dating from 40.0 to 12.7 ka in Kings
Canyon (920–1,270 m; Cole 1983), record a mesic mixedconifer forest consisting of Pinus cf. ponderosa, P. lambertiana, Abies grandis, Calocedrus decurrens, Juniperus occidentalis and Pinus monophylla, with an understory of Ceanothus
integerrimus, Garrya flavescens, and a single record of Torreya
californica. Pollen of Sequoiadendron giganteum suggests a
stand close to the site. Beginning at 13.7 ka, pollen sites
record increasing conifer pollen and macrofossils (Abies concolor, Pinus lambertiana, Pinus ponderosa, Calocedrus decurrens, and Sequoiadendron) suggest that mixed-conifer and
subalpine forest ascended the west slope at least to the altitudes of these pollen sites. The transition from steppe to forest took place in 10.0 ka in the Stanaslaus River drainage of
the northern Sierra Nevada.
Late glacial maximum macrofossil assemblages along
coastal California (40.0–14.0 ka) at Tomales Bay, Carpinteria,
Santa Cruz Island, and Rancho La Brea Tarpits (summarized
in Johnson 1977) record floras comparable to vegetation seen
today, but with a southward extension of some mesic species.
These sites record chaparral species in Adenostoma, Arctostaphylos, Ceanothus, Garrya, Heteromeles, Quercus, and
Rhamnus and closed-cone conifer forests of Pinus radiata and
P. muricata. Closed-cone conifer forests were also recorded at
Point Sal Ridge and San Miguel Island. The Tomales Bay flora
records riparian forests (Acer macrophyllum, Alnus rubra),
mixed-evergreen forest (Arbutus menziessii, Pseudotsuga menziesii, Umbellularia californica, Quercus agrifolia, Torrya californica) as well as Picea sitchensis, now found as far south as Cape
Mendocino. Chaparral at Carpenteria (Arctostaphylos, Ceanothus) contained trees now common in the area including
Quercus agrifolia, Umbellularia californica, and Pseudotsuga
58
macrocarpa. The Carpinteria flora also contains mesic and
xeric trees now exotic to this region including P. muricata, P.
radiata, Cupressus forbesii, C. goveniana, C. macrocarpa, Juniperus californica, Pinus sabiniana, and Sequoia sempervirons.
Nearby Santa Cruz Island has late glacial records in Arctostaphylos, Ceanothus, and Garrya and Pinus remorata (P.
muricata), as well as the exotic Pseudotsuga menziesii and
Cupressus governiana. The Rancho La Brea site in Los Angeles
has a rich assemblage of chaparral species now common in
the area (e.g., Adenostoma fasciculatum, Arctostaphylos glauca)
or found on the channel islands and central California coast
(Arctostaphylos viscata, A. morroensis). Modern tree elements
include Alnus rhombifolia, Juglans californica, Platanus racemosa, Quercus agrifolia, Q. lobata, as well as exotic species
Pinus muricata, P. radiata, P. sabiniana, Cupressus forbesii, C.
goveniana, C. macrocarpa, Juniperus californica, and the mixedevergreen forest component Sequoia sempervirons. The McKittrick flora in the southern San Joaquin valley contains
pinyon-juniper woodland of Pinus monophylla and Juniperus
californica with Arctostaphylos glauca and A. pungens. Pinyonjuniper now occurs 50 km from this site at the south end of
Cuyama Valley.
Packrat middens in the Mojave and Sonoran Deserts
(Spaulding 1990; Van Devender 1990; McAuliffe and Van
Devender 1998) show a rich late Pleistocene-Holocene fossil
flora. Because fossil middens are selectively preserved in rock
shelters, midden vegetation records tend to be biased to
mountainous slopes rather than alluvial bajadas and plains.
In the late Pleistocene (25–12 ka), Mojave Desert middens
show that pinyon-juniper woodlands formed extensive cover
from 1,000 to 1,800 m, with juniper woodlands occurring 1,000 m elevation. Only the lowest basins such as Death Valley were free of tree cover (Wells and Woodcock 1985; Woodcock 1986). Associated species in pinyon-juniper woodlands
include arboreal leaf succulents Yucca schidigera, Y. brevifolia,
Y. baccata, and Y. whipplei and subshrub understory consisting of species in Artemisia, Chrysothamnus, Lycium, Ephedra,
and Atriplex. The highest peaks, such as those in the Clark
Mountains, had subalpine forests of Pinus longaeva and P. flexilis. In the Owens Valley, middens at Owens Lake and the
Alabama Hills record Utah Juniper-pinyon woodland with
Yucca brevifolia and understory of Great Basin sagebrush and
Mojave Desert scrub including Artemisia tridentata, Purshia tridentata, Atriplex confertifolia, Ericameria cuneata, Ephedra
viridis, and Glossopetalon spinescens (Koehler and Anderson
1994, 1995). A pollen record from Owens Lake (Mensing
2001) reveals that juniper woodland dominated Owens Valley in the period 16.2–15.5 BP (years before present). Woodlands were replaced by shrubland (mostly sagebrush) in
15.5–13.1 (cal) BP. Chenopodiaceae soon replaced Artemisia.
Pinyon-juniper woodland (Pinus monophylla, Juniperus californica), desert chaparral (Quercus turbinella, Ceanothus greggii),
and Mojave Desert scrub were recorded in the Scobie Mountains on the eastern escarpment of the Sierra Nevada
(McCarten and Van Devender 1988). Middens in the White
Mountains and nearby Volcanic Tablelands at 1,300 m record
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Utah juniper with understory of desert scrub of Purshia tridentata, Tetradymia axillaris, T. canascens, and Ericameria
cuneata (Jennings and Elliott-Fisk 1993).
In the Sonoran Desert (Van Devender 1990), middens
record pinyon-juniper woodlands with scrub live oak (Quercus turbinella, Q. dunnii) to as low as 500-m elevation. There
were few modern Sonoran Desert plants at this time. The
lowest basins from 600 to 300 m had juniper woodlandshrub live oak assemblages with Juniperus californica, Yucca
brevifolia, Y. whipplei and Nolina bigelovii. Most of these
species grew 400 to 1,200 m below present elevations. Members of Creosote bush scrub, including the modern dominants Larrea tridentata, Ambrosia dumosa, and Encelia farinosa were present in full-glacial climate in the Tinajas Atlas
of the upper Gulf of California at 18 ka. Middens show that
wildflowers in the same genera were present in the Mojave
Desert since at least the late Pleistocene (Table 2.3). It
appears that there has been little change in the distribution
and composition of desert annuals, even though there were
large changes in woody cover. Because annual forb species
persist on seed banks, selection may have been lacking for
long-range seed dispersal, even with climate change as large
as that in the Pleistocene-Holocene transition. Conifer
woodlands and chaparral extended as far south as central
Baja California. Juniperus californica and chaparral species
were recorded in the Sierra San Francisco at 10.2 ka (Rhode
2002). J. californica and Pinus quadrifolia and chaparral
(Adenostoma fasciculatum, Quercus turbinella) were recorded
in a full-glacial midden (17.5 ka) in the northern Vizcaino
Desert. J. californica, Pinus quadrifolia, and chaparral dominated by Adenostoma fasciculatum and Quercus turbinella persisted on the site until at least 10.0 ka (Wells 2000).
EAR LY AN D M I DDLE HOLO C E N E
The Pleistocene-Holocene transition at ca. 10 ka was a period
of profound global climate change with deglaciation. The
Laurentide ice-sheet collapsed by ca. 8,000 BP (Dawson 1992),
and Milankovitch solar maximum appears to have contributed to warmer and drier climates than at present. GCM
simulations by Bartlein et al. (1998) suggest that the modern
global circulation was fully established in the western United
States. However, the combination of remnant ice sheets over
northeastern Canada (Dyke and Prest 1987; Broecker et al.
1989), colder continental winters (the Earth’s orbit was at perihelion), and higher Pacific Ocean SSTs may have encouraged
the poleward displacement of moist surface westerlies to
British Columbia (Bartlein et al.). Direct thermal evidence for
warmer temperature comes from bristlecone pines (Pinus longaeva) that were analyzed for stable hydrogen isotopic composition (Feng and Epstein 1994). Based on ring chronologies
matching living trees with dead trees in a continuous timeseries from 8,000 to present, they found that temperatures
reached maximum levels at 6,800 BP. Submerged pine stumps,
dated 6.3–4.8 ka (cal), revealed that Lake Tahoe was below sill
height (Lindstrom 1990). Nearby Walker Lake, which drains
the eastern Sierra Nevada south of Lake Tahoe, was shallow
and desiccated from 5,300–4,800 BP (Benson, Meyers, and
Spencer 1991). Clark and Gillespie (1997) provide evidence
that glaciers in the Sierra Nevada may have been entirely
absent during much of the Holocene. Konrad and Clark
(1998) obtained three cores from a meadow directly below the
Powell Rock Glacier and found only nonglacial grus from the
early Holocene. Evidence of early Holocene aridity is also seen
in the West Berkeley shell mounds (Ingram 1998). Radiocarbon ages of marine shells and charcoal suggest changes in the
14 C content of San Francisco Bay surface waters relative to
the atmosphere (the oceanic reservoir age) was high at the
beginning of the record at 4,970 BP. This period is coincident
with relatively high salinity in San Francisco Bay and dry climate throughout California. Changes in the radiocarbon
reservoir age may be due to changes in the strength of seasonal wind-driven upwelling off coastal California, where
upwelling brings 14C depleted waters to the surface. Ingram
proposes that the region was influenced by an abnormally
strong Pacific anticyclone, northward displacement of the jet
stream, and low rainfall. Enhanced upwelling is physically
consistent with the poleward displacement of moist westerlies
along the Pacific coast (Bartlein et al. 1998).
Pollen evidence in the Oregon Cascade Range shows
open conifer forests of Pseudotsuga and Quercus with a
Pteridium understory in areas now covered with mesic taxa
in Thuja, Tsuga, and Picea (Thompson et al. 1993; Sea and
Whitlock 1995; Long et al. 1998). At Diamond Pond in the
Great Basin, shadescale expanded at the expense of sagebrush during 7,000–4,000 BP (Wigand 1987; Thompson
1990). In the central Sierra Nevada, pollen and macrofossils
at localities such as Giant Forest and Starkweather Pond
suggest that forests had ascended to current limits by
10.5–9.0 ka. Some conifers were still colonizing new habitat as late as 6.3 ka (Anderson 1990). Pollen records also
show a decrease in Pinus and Abies, and an increase in
Cupressaceae and Quercus (Woolfenden 1996). Increasing
charcoal suggests increasing fuel buildup and densification
of stands. However, Anderson (1990) indicates that forests
were more open and contained abundant shrubs, such as
Chrysolepis sempervirens, compared to forests at present.
Pinus contorta experienced successive disappearance from
lower elevation sites (Anderson 1996). At Nichols meadow,
pollen evidence documents the development of a lower
montane forest and giant sequoia (Sequoiadendron giganteum) community around a small meadow between 12.5
and 8.8 ka (Koehler 1994). Poor pollen preservation 8.8–6.0
BP was attributed to early Holocene aridity. Barrett Lake at
2,900 m was colonized by forest by 10.0 ka, with shrub
pollen increasing from 10.0 to 6.0 ka (Anderson 1990).
Scuderi (1987a) found that the Sierra Nevada tree line at
Cirque Peak was 70 m higher than at present, with the
maximum levels occurring between 7,150 and 3,800 (cal)
BP. The tree line was 125 m higher in the White Mountains
(La Marche 1973). These displacements reflect a global
trend, as higher tree lines were reported in Scandinavia,
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TA B L E 2.3
Fossil Wildflowers and Bunch Grasses in Packrat Midden in the Mojave and Sonoran Deserts
Late
Pleistocene
Pleistocene
Species
Grand Canyon
Marble Mtns.,
Owl Canyon,
Point of Rocks
Mead and
Phillips (1981)
Spaulding
(1983)
Amsinckia
✓
✓
Argemone
✓
✓
Late PleistoceneHolocene
Lucerne Valley
Lower Colorado
River Valley
Greenwater
Valley
Cole (1986)
King (1976)
King and
Van Devender
(1977)
Cole and
Webb (1985)
✓
✓
✓
✓
✓
✓
Pichaco Peak
✓
Aristida
✓
Astragalus
✓
✓
✓
Castilleja
✓
Chorizanthe
Circium
✓
✓
✓
✓
Cryptantha
✓
✓
✓
✓
Draba
✓
Dithyrea
✓
Eriogonum
Eschschozia
✓
✓
✓
✓
✓
✓
✓
✓
Hilaria
✓
✓
✓
✓
Lupinus
✓
✓
Malvastrum
Mentzelia
✓
Mirabolis
✓
✓
✓
Oryzopsis
✓
Penstemon
✓
✓
Perityle
✓
Pectocarya
Plantago
✓
✓
Gilia
Phacelia
✓
✓
Euphorbia
Lepidium
✓
✓
✓
✓
✓
✓
✓
Plagiobothrys
✓
✓
Solanum
Stephanomeria
Stipa (Nessela)
Vulpia
Other grasses
Late Holocene
✓
✓
✓
✓
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Carpathian Mountains, European Alps, and the Rocky
Mountains (Bradley 1999).
Packrat middens show that pinyon-juniper woodlands in
the Mojave Desert in full-glacial climates were quickly
replaced by modern desert scrub assemblages in the early
Holocene (Spaulding 1990; Van Devender 1990). In the
Sonoran Desert, Pinus monophylla dropped out by 11.0 ka,
leaving live oak woodland or chaparral with Juniperus californica and Lycium specoes continuing to 9.0 ka. Pinyon-juniper
retreated upslope to higher ranges exceeding 1,200 m. Creosote bush scrub (CBS) was recorded as far north as Picacho
Peak near Yuma Arizona by 12.7 ka and the Whipple Mountains near the Colorado River at 10 ka. By 10 ka, much of the
Mojave Desert was covered with desert scrub but the flora still
lacked Larrea tridentata and other members of CBS. However,
Cole (1986) indicates that CBS in the lower Colorado River
Valley had a modern aspect beginning at 12 ka. In the
Mojave Desert middens record L. tridentata, Encelia farinosa,
and Ambrosia dumosa at the Marble Mountains by 7.9 ka
(Spaulding). At Lucerne Valley a high desert flora recorded
between 11.0 and 7.8 ka contained A. dumosa. Modern CBS
was recorded in middens dating to 5.9–3.6 ka (King 1976). L.
tridentata reached the Skeleton Hills northwest of Las Vegas
by 8.2 ka and Ambrosia dumosa was recorded at Death Valley
by 10 ka. Owens Lake records indicate a decrease in Juniperus
pollen. Packrat middens in the Alabama Hills record the
arrival of Chrysothamnus teretifolius, Grayia spinosa, Lycium
andersonii, and Krascheninnikovia lanata (Koehler and Anderson 1995). Middens dated 8.7–7.8 ka in the White Mountains
at 1,830 m, record pinyon-juniper-sagebrush woodland with
a similar species composition to that at present (Jennings and
Elliott-Fisk 1993). This record appears to capture the northward expansion of Pinus monophylla because a 9.8 ka midden
in the nearby Volcanic Tablelands did not record the species.
The full glacial northern limit was at the southern Inyo
Mountains (Koehler and Anderson 1994). Consistent with
higher tree lines in the White Mountains and Sierra Nevada,
a midden at 5.6 ka at 3,038 m in Silver Canyon of the White
Mountains records P. monophylla and Juniperus osteosperma at
higher elevations than at present, in areas now covered with
subalpine forest of Pinus longaeva and P. flexilis. Packrat middens in northwestern Nevada reveal only small changes in
pinyon-juniper and shrubland elements, even after the Pleistocene-Holocene transition (Nowak et al. 1994).
From genetic evidence, Ledig (2000) hypothesized that P.
coulteri and P. sabiniana, both members of Pinus sect. sabinianae, migrated northward in postglacial times. He found
that the genetic heterozygosity in Pinus coulteri decreased
with latitude, with the decrease paralleled by a loss in alleles north of the Peninsular Ranges at points with large gaps
in its distribution. He hypothesized that the genetic pattern
reflects a cascading series of founder events as P. coulteri
invaded the Transverse and South Coast ranges from Pleistocene sites farther south. Alternatively, the high allelic
diversity of southern California Coulter pine populations
may be due to in situ genetic drift due to periodic isolation
caused by extensive mortality and range contraction from
extreme drought, such as that that experienced in 2001–2002.
Similarly, Ledig (1999) found a decrease in heterozygosity of
P. sabiniana decreases northward in California, but genetic
differences between populations on both sides of the Great
Valley are small, similar to that in continuous populations.
The fossil record of P. sabiniana is restricted to southern California during the late Miocene and Pliocene. He concluded
that P. sabiniana formed a continuous forest in the Central
Valley in the late Pleistocene, with its range becoming fragmented in the early Holocene.
The early Holocene may have been wetter than present in
far southeastern California. Wells and McFadden (1987)
found several deep lakes in the Lake Mojave basin between
9.5 and 8 ka, and attributed it to enhanced summer monsoon of the Milankovich solar maximum. This is supported
by late Holocene lake stands that persisted until 7,450 BP at
Laguna Chapala in central Baja California (Davis 2003). An
intensified summer monsoon, rather than winter rains, is
the most likely explanation because the jet stream had
already retreated to modern latitudes by this time, far north
of Laguna Chapala. Tropical cyclones reaching the southwestern United States may have produced heavy rains more
frequently compared to the present. The persistence of
Juniperus californica and chaparral species in the California
and Baja California deserts into the early Holocene (Wells
and McFadden 1987; Wells 2000; Rhode 2002) may reflect
phenological plasticity from winter to summer rains, until
the decline of summer rains resulted in extirpations by ca.
7,000 BP.
M I D -HOLO C E N E
The mid-Holocene in California (5.0–3.5 ka) is a period of
greater rainfall, enlargement of Sierra Nevada alpine glaciers
and small interior lakes. The most important trend in GCM
simulations is the increasing presence of moist westerlies on
the Pacific Coast (Bartlein et al. 1998). West Berkeley shell
mounds (Ingram 1998) show that the 14C content of San
Francisco Bay surface waters has low radiocarbon reservoir age
from 4,970 to 2,000 BP and especially from 3,800 to 2,800 cal
yr BP. This trend was attributed to a reduced strength of seasonal wind-driven upwelling off coastal California. This
period is also coincident with relatively low salinity in San
Francisco Bay (indicating high freshwater inflow) and wet climate in California. Benson, Meyers, and Spencer (1991) found
that Walker Lake was relatively high between 5,100 and 3,170
(cal) BP. High stands occurred between 4,000 and 3,600
(uncalibrated) BP at Mono Lake, Silver Lake, and Searles Lake
(Stine 1990; Enzel et al. 1992). Peat deposits developed in this
period at Ash Meadows in the Armagosa Desert. Analysis of
marine and nonmarine groups in a pollen core at Newport
Bay, southern California, show periodic intrusions of fresh
habitat from 4,190 to 2,500 (cal) BP (Davis 1992).
In the Sierra Nevada, glaciers established in areas where
they were apparently absent in the early Holocene. The
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Recess Peak advance in alpine cirque glaciers in the central
Sierra Nevada is based on morainal deposits aged at 3,000
BP on the basis of lichemitry (Burke and Birkeland 1983). A
lake core from below Conness glacier reveals alternating
gyttja and silt after about 3,900 (cal) BP (Konrad and Clark
1998). The transition from higher organic sediments to
lower organic sediments, at 3,400 (cal) BP indicate formation and growth of the glacier in the cirque at that time.
Cosmogenic records of moraines left by small headwall glaciers on Mt. San Gorgonio give Holocene dates, but the resolution of the method does not permit correlation with the
mid-Holocene advance in the Sierra Nevada (Owens et al.
2003).
Several localities reveal a shift to more mesic vegetation.
In the Cascade Mountains there was a decrease in Quercus
and Pseudotsuga, and an increase in Thuja, Tsuga heterophylla, Picea, and Alnus rubra (6,850–2,750 BP; Long et al.
1998; Thompson et al. 1993). At Diamond Pond in the
Great Basin, sagebrush gradually replaced shadescale
between 5,400 and 4,000 bp, with juniper and grasses reaching greatest abundance at 4,000–2,000 BP. Pollen profiles
show short-term maxima in juniper abundance at 3,500
and 2,500 BP and a grass/sage maximum at 3,600 and 3,100
BP (Wigand 1987; Miller and Wigand 1994). There were few
changes in the floristic composition of Great Basin woodlands during the late Holocene, although pollen evidence
suggests that sagebrush expanded and shadescale steppe
contracted (Thompson 1988, 1990). Fossil middens show
little vegetation change in the Mojave Desert, although
middens in the Sheep Range of Nevada record a shift from
pinyon dominance at 3,520 BP to juniper dominance at
1,990 BP (Spaulding 1985, 1990). Middens in northwestern
Nevada show that Pinus monophylla entered this region only
by 3,000 BP (Nowak et al. 1994). This trend is attributed to
a postglacial migrational lag rather than climate change due
to frost and soil heaving, as evidenced by the absence of
P. monophylla from cold basin floors. P. monophylla recruitment appears to be dependent on nurse plants until canopy
closure (Wangler and Minnich 1996; Chambers 2001).
Hence, P. monophylla migration may have been limited to
contagious diffusion of mature stands rather than more efficient long-distance disjunct expansions (see Clark 1998). It
is proposed that fires carried by dense fields of wildflowers
in wet years occurred more frequently than in other periods
of the Holocene (Minnich et al. in review).
Pollen abundances on the midslopes of the Sierra Nevada
show an increase in Abies and Cupressaceae (Calocedrus
decurrens) relative to Pinus and a decrease in Quercus (Davis
and Moratto 1988; Anderson 1990). Pollen profiles along
the crest of the Sierra Nevada record an increase in upper
mixed-conifer and subalpine forest species, especially Tsuga
mertensiana and Abies magnifica by 6,000 BP (Anderson
1990). Pollen and macrofossil data suggest that Pinus contorta experienced a return to lower elevation sites by 1,700
BP. Early Holocene tree lines in the southern Sierra Nevada
persisted until 3,500 BP, then declined in response to cooler,
62
wetter conditions (Scuderi 1987a; Lloyd and Graumlich
1997). Scuderi argued that tree lines declined about 70 m at
3,400 and 2,400 BP, similar to trends in the White Mountains (La Marche 1973). He attributed the decline to Feng
and Epstein’s (1994) decreasing temperature profile
recorded in White Mountain bristlecone pines. Graumlich
(1993) proposed that tree line stands experienced no
recruitment and that old trees slowly perished during marginal episodes.
Middens in the Sonoran Desert show that desert microphyll woodland species such as Cercidium floridum and
Olneya tesota did not arrive into the region from Mexico
until 6,000–4,000 BP (Van Devender 1990; McAuliffe and
Van Devender 1998). These studies suggest that desert
microphyll species are frost-sensitive and were kept out of
the region by colder winters caused by a Milankovitch winter insolation minimum of the early Holocene.
LATE HOLO C E N E DROUG HT
An extended period of reduced precipitation existed in
2,000–600 BP that culminated in profound drought beginning 1,000 BP. Ingram (1998) found the highest C-14 content of marine shells over the past 5,000 years at
1,900–1,200 (cal) BP. She pointed out that the age of the
top of the West Berkeley mound and several other mounds
in the San Francisco Bay region (1,100–1,300 [cal] BP) coincided with a prolonged dry period in California, stronger
northerly winds, coastal upwelling, and low river inflow to
San Francisco Bay. Evidence for regional drought is seen in
low lake stands at Walker Lake and Mono Lake (Stine 1990;
Benson 1991). Tree stumps rooted in present-day lakes,
marshes, and streams of the Sierra Nevada were attributed
to severe drought 1,000–700 BP (Stine 1994). Scuderi
(1987b) found evidence that tree-ring widths in Pinus balfouriana were inversely related to glacial expansions, so he
was able to deduce (from wide rings) that a period of pronounced warmth had existed during 1,200–1,000 BP.
Packrat midden data are too sparse to indicate broad scale
effects of the Medieval Drought on desert vegetation.
Spaulding (1990) noted small movements of vegetation
zones upslope by as much as 100–200 m between 1,500 and
500 bp, and Van Devender (1990) found little evidence of
vegetation change. At Greenwater Valley in the eastern
Mojave Desert, Cole and Webb (1985) described a decrease
in Larrea tridentata since 2,235 BP, but most other shrubs
remained unchanged during the period. Sparse fossil charcoal at El Paso Peaks seismic trench suggests limited fire
activity between 2,800 and 700 (cal) BP.
LIT TLE IC E AG E
The “Little Ice Age” (700–200 BP) was a global event that
caused the expansion of mountain glaciers in Europe, North
America, and elsewhere (Lamb 1995). In the Sierra Nevada,
fresh moraines (virtually lacking lichen) of the little ice age
Matthes advance are found below hundreds of modern and
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extinct cirque glaciers (Burke and Birkeland 1983; Clark and
Gillespie 1997). Hydrogen stable isotope data for bristlecone
pine in the White Mountains indicate a period of rapid cooling in 1,600 CE, with temperatures during the remainder of
the Little Ice Age being as cold as any period in the Holocene
(Feng and Epstein 1994). The Little Ice Age is associated with
high lake stands between 600 and 300 bp at Mono Lake, Silver Lake, and Cronise Lakes at the mouth of the Mojave
River, as well as episodes of marsh expansion in the
Panamint Valley and peat formation in Ash Meadows (summarized in Enzel et al. 1992). Pulses of sediment activity
developed in the Kelso dunes, apparently derived from the
Mojave River date to 690 and 450 BP (Clarke and Rendell
1998). There was another intrusion of fresh water habitat at
Newport Bay at 560 BP (Davis 1992). Scuderi (1990) found a
sharp decline in indexed ring width of timberline trees at
Cirque Peak at 1600 CE. Tree ring widths were inverse to glacial expansions in Pinus balfouriana, with minima in widths
in the years 810, 1470, 1610, 1700, and 1810 CE. Graumlich
(1993) found a period of cold temperatures from ca. 1450 to
1850. Lloyd and Graumlich (1997) described a decline in
treeline and treeline forest abundance from 450 to 50 BP.
Tree ring data for Pseudotsuga macrocarpa indicate more frequent wet winters in late 1500s and 1600s than in recent
centuries (Michaelson and Haston 1988).
There is limited evidence of vegetation change during the
Little Ice Age. Wigand (1987) found that juniper abundance
increased at Diamond Pond at 300–150 bp. At Greenwater
Valley in the eastern Mojave Desert, there was a small downward shift in Coleogyne ramosissima of 100 m between 205
and 565 BP (Cole and Webb 1985). Examination of vegetation descriptions in late 18th-century Spanish diaries in
northern Baja California show the broad scale distribution,
species composition, and local patterning of communities
are consistent with those today (Minnich and Franco 1998).
Summary of Tertiary–Quarternary Climate
and Vegetation Change
Vegetation lifeforms of California and adjoining regions
experienced different fates over the Cenozoic as the climate
gradually cooled and precipitation shifted from year-round
to a winter-wet mediterranean regime. As winters grew
colder, ancestral Eocene conifer forests along the Rocky
Mountains moved to nearby Great Basin terrains and established in the rapidly rising California mountain ranges by the
late Miocene-Pliocene. Conifer forests were favored in part by
the increasing presence of the westerlies and winter precipitation. Drought-tolerant conifers also benefited from colder
climates because they could adapt to developing summer
drought with evergreen sclerophyll metabolism. A few genera
(Ginkgo and Metasequoia) that apparently had high moisture
demands, became extinct in the western United States. In the
Great Basin, montane conifer forest was replaced by pinyonjuniper-chaparral woodlands in the Pliocene due to rain
shadows created by the uplift of the Sierra Nevada. Ancestral
mesosphytic deciduous forests with alliances to modern floras in east Asia and the southeastern United States also
moved westward from the Rockies, but they were selectively
eliminated by developing summer drought; their survivors
now occupy riparian and aquatic habitats. While global cooling appears to have eliminated Tertiary tropical floras in California, ancestral temperate evergreen sclerophyll shrubs and
trees first recorded in the southern Rocky Mountains, gave
rise to descendents that reached California in the Oligocene
and Miocene. These elements now comprise the modern
mixed evergreen forest, oak-laurel woodland, and chaparral.
A drying climate eliminated such woodland components as
Persea and Ocotea, but others that had been recorded in
Pliocene Nevada still survive on southern California islands
(Lyonothamnus, Prunus lyonii) and in the Santa Lucia Mountains (Abies bractiata).
Although a fossil record is lacking, the presence of high
altitude accretionary terrains and later extensional terrains
in Mexico suggests that temperate, seasonally dry subtropical climates there may have contributed to the diversification of evergreen sclerophyll floras in such genera as Pinus,
Quercus, and chaparral genera through most of the Tertiary.
(In this view, the California flora is a northern extension of
a Mexican flora.) Similarly, the relationships between paleofloras and the uplift of the Sierra Nevada and Great Basin
are poorly constrained. This is especially critical in explaining the rich paleofloras in the Great Basin during the
Miocene. The displacement of Rocky Mountain forest by
pinyon-juniper woodland could be related to either the subsidence of the Great Basin or the Mio-Pliocene uplift of the
Sierra Nevada and southern California Coast Ranges; both
would produce warmer climates and/or reduce precipitation. The Miocene decline in Pacific SSTs due to Antarctic
glaciation would enhance rain shadowing by reducing convective precipitation and increasing the role of physiographic lift in winter frontal storms that increasingly dominated total precipitation.
Geophysical models of a continuously high Sierra Nevada/
Great Basin since the Cretaceous must account for the presence of subtropical chaparral, oak-laurel floras, and pinyonjuniper woodlands in the Sierra and Great Basin in the
Miocene. Under standard atmospheric lapse rates, climates
at 3 km elevation have supported freezing temperatures
either in winter or at night throughout geologic time, and
oxygen isotope ratio curves show that global temperatures
were approaching modern levels by the Miocene. It is
unlikely the subtropical floras could have survived as high as
2 km.
Axelrod (1973) proposed that ancestral chaparral was
seral to oak-pine woodlands in summer rain climates in
Mexico. California chaparral replaced woodland only by the
Pleistocene, due to the rapid uplift of the Pacific Coast
ranges, with widespread expansion in the late Holocene.
This interpretation is narrow because data on paleoterrains
and the findings of GCMs are poorly constrained. All the
ingredients for a mediterranean climate were in place by the
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end of the Miocene, in particular cold upwelling currents
contributing to summer drought. Fires are frequent in modern Mexican pine-oak woodlands and chaparral, especially
in late spring, before the arrival of the North American
monsoon. Some Mexican mountain ranges (e.g., Sierra
Madre Occidental) today support extensive chaparral in a
summer-wet climate (Vankat 1988; Rzedowski 1978). There
is insufficient evidence to determine whether chaparral
became extensive 10 ka or 10 ma in California.
The paleovegetation history in the western United States
shows that individual species each have independent histories of evolution and migration. These are inseparable
processes because selection operates in coupled spatial and
temporal scales (Davis and Botkin 1985; Betancourt et al.
1990; Sauer 1988), resulting in temporally loose, changeable
associations. For example, early Tertiary Rocky Mountain
conifer forests grew with deciduous hardwoods but these
have since parted ways. Chaparral understory in Pliocene
pinyon-juniper woodlands of Nevada were replaced by
desert scrub as climates became drier and cooler. Packrat
middens show that desert woodland species experienced
divergent shifts in elevational zonation in response to Pleistocene-Holocene climatic change (Spaulding 1990; Van
Devender 1990).
At the scale of life forms, the evolution and migration of
species appear to be clustered geographically; that is, ecological convergence, because plants with similar metabolisms often find comparable paths of adaptation within a
climatic range. In particular, the California chaparral has
been the object of extensive studies of convergence
(Mooney 1977), but what is the nature of convergence over
long time scales? The flora at any locality represents the
cumulative effect of in situ evolution and species immigration. The importance of each depends on the pace of evolution versus the pace of climatic change (Davis and Shaw
2001). The relationship between convergence in chaparral
and climatic change is seen in two end-member evolutionary models (Minnich 1985; Barbour and Minnich 1990). In
one model the species is not initially well adapted to the
new environment. Thus species are continuously shaped by
changing climate, and climatic change shapes genetic and
species changes. The other end-member model is a generalist species that is well adapted to a wide range of environments, and performs in the same adaptive mode over long
spans of geologic time. It is adapted to climate in a general
way with no experimentation or speciation. As a new climate develops, the conservative generalist species expand
into the new environment without evolutionary change
(migration without evolution). The California chaparral
appears to be a combination of these models.
The Tertiary record indicates that the ancestors of chaparral
predate mediterranean climate in the southern Rockies and
spread westward to California by the Miocene. Axelrod (1973)
suggests that many evergreen sclerophyllous shrubs are similar to species in Miocene records, despite opportunities for
evolution. Evidence for the generalist model is seen in mod-
64
ern biogeography. Most California chaparral genera have few
often distantly related spp (eg. Rhamnus), one species
(Anorthostaphylos), or are monotypic genera (Heteromeles, Malosma). Most shrubs live in summer rain climates in Arizona,
central and southern Baja California, and mainland Mexico.
The range of chaparral genera into Mexico suggests that evergreen sclerophyll metabolism is a generalized adaptation to
drought, regardless of seasonality (Valiente-Banuet et al.
1998). Many chaparral genera may date to the Eocene while
mediterranean climate dates only to ca. 10 ma. Perhaps the
clearest case of recent evolution is the high diversity of Arctostaphylos and Ceanothus, but the radiation of these genera
may be more related to selective pressures produced by fire
than drought adaptation. Most species are nonsprouters—an
unusual trait in angiosperms—with obligate seeding increasing the magnitude of evolutionary change. The chaparral
appears to be a mixture of evolutionary convergence and
migrational convergence, the former being evolution and
range expansion within a climate, and the latter being a geographic expansion of similar adaptive forms without evolution into a newly developing mediterranean climate.
California ecosystems experienced large shifts in the latitudinal and altitudinal zonation of vegetation since the last
glacial maximum. Perhaps the clearest example is that Sierra
Nevada subalpine forests now occupy extensive areas that
were covered with mountain ice sheets in the late glacial
maximum. The available paleobotanical data permit only
broadscale reconstruction of vegetation assemblages. The
strongest case has been made from packrat midden records
on the replacement of late glacial pinyon-juniper woodlands by creosote bush scrub in the Mojave and Sonoran
Deserts (Spaulding 1990; Van Devender 1990).
There is little consensus on the character of full-glacial climates in California. Using climatic transfer functions, temperatures at Clear Lake at northern California in 18 ka were
estimated to be 8C cooler than present (Adam 1981; Adam
and West 1983). Precipitation was estimated to be an astonishing 200 cm greater than the present AAP of 120 cm. However, the lapse rate method used in the study overestimates
full-glacial precipitation and temperature changes because it
assumes that vertical lapse rates account for latitudinal displacement of temperature and precipitation fields. A 7 C–8C
decrease in mean annual temperature is physically inconsistent with estimates of 1 C–3 C decreases in Pacific Ocean SSTs
(CLIMAP 1976; Moore et al. 1980; Lowe and Walker 1998).
Modern mean winter temperature of 5 C at Clear Lake suggests that full-glacial temperatures averaged –2C to –3C.
Hence, average atmospheric lapse rates would require the
Pacific Ocean SSTs to be near freezing, but this is unsupported
by oxygen isotope records from deep-sea cores (Bradley
1999). Using pollen records in the Sierra Nevada, Koehler
(1994) concluded that high Artemisia and Poaceae pollen percentages for 18.5–12.5 ka was evidence for a cold-dry steppewoodland environment with low snow melt rates. However,
wind-pollinated dominants such as Artemisia tridentata
presently grow at altitudes ranging from 600 to 3,800 m, thus
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limiting the climatic indicator value of the species. A dry climate is inconsistent with the year-round presence of a strong
jet stream. Climates far colder than at present are unlikely
because ocean temperatures were only 1C–4C below those
at present (Shackelton 1987; Bartlein et al 1998; Kaser and
Osmaston 2002). Solar radiant energy, the dominant energy
source in snowmelt (Miller 1981), was as high in the late glacial as at present (Bradley 1999). Hence, Sierra Nevada ice
sheets are evidence for heavier snowfall, not reduced
snowmelt. GCMs of Bartlein et al. suggest that mean annual
precipitation was perhaps 150%–200% of modern levels.
Small decreases in winter temperatures were paralleled by
large decreases in summer temperatures [thermal climates
became more equable], especially in the mountains and interior valleys because the Laurentide ice sheet maintained a
strong jet stream over the region. Resulting Pacific air intrusions prevented the buildup of hot summer air masses seen
today over the western United States. Stable air masses of the
coastal marine layer may have maintained summer drought
west of the California mountains, but light precipitation may
have occurred in the Great Basin, perhaps comparable to
modern late-spring climate.
Low summer temperatures suppressed productivity and
tree lines, much as seen in mediterranean climates in the
Andies of central Chile where the combination of a strong jet
stream supported by the Antarctic ice sheet and smaller continental area produce summer temperatures 5C–8C cooler
than at comparable latitudes of California. The modern tree
line lies near 2,000 m, a full 1,000 m lower than in California. A moist Sierra Nevada climate appears to be reflected in
full-glacial middens in Kings Canyon which records a rich
mixed-conifer forest in an area now cover by oak-pine-chaparral woodland (Cole 1983). Pleistocene middens along the
California give records of mixed evergreen forest, closedcone conifer forest, and chaparral, with the southward displacement of Picea sitchenisis, Sequoia sempervirons, and
Pseudotsuga menziesii. The persistence of summer drought,
predicted by GCM simulations, is supported by the presence
of chaparral at several sites in central and southern California. Summer drought very likely maintained fire regimes
comparable to those at present, only the distribution of
burning was cued to full-glacial distributions of flammable
herblands, chaparral and conifer forests. In the interior, rain
shadows from the California coastal ranges, Transverse and
Peninsular Ranges, and the Sierra Nevada resulted in semiarid climates, as evidenced by the presence of sagebrush and
Mojave Desert scrub assemblages in most midden sites.
Sierra Nevada montane forests species did not reach the
base of the eastern escarpment (McCarten and Van Devender 1988). Cooler temperatures and reduced evapotranspiration contributed to the buildup of the pluvial lakes that
were maintained by runoff from the Sierra Nevada and
other high ranges. Summer precipitation may have been
lower than present in the Mojave and Sonoran Deserts
because the stronger westerlies displaced the North American monsoon into Mexico.
Rapid vegetation change occurred in response to climatic
change of the Pleistocene-Holocene transition. Members of
creosote bush scrub advanced rapidly across the Sonoran
and Mojave Deserts, reaching their modern limits by ca. 8
ka. Moreover, plant migrations appeared to have lagged
behind climatic fluctuations, even at scales of millenia. In
long-lived communities, a span of centuries is brief for even
short-term population changes. Selection processes that
result in changes in recruitment or disturbance patterns
require several generations to reach mature phases of vegetation (Sauer 1988; Thompson 1988, 1990). Some species
were undergoing range adjustments in the late Holocene.
Pinus monophylla may have reached the Reno area only in
the past 2.0 ka (Nowak et al. 1994). In the central Sierra
Nevada, Anderson (1990) concluded that open forests at 10
ka reflected drier conditions than at present. However, the
openness of stands may also reflect poor soil development
after deglaciation. Van Devender (1990) concluded that
desert scrub changed little since the early Holocene suggesting that modern floras have equilibrated to modern climate.
The paleobotanical record for the mid- and late Holocene
reveals limited vegetation change under strong vegetation
inertia. Climatic shifts in precipitation as evidenced by small
shifts in lake stands and glacial advances appear to have
been insufficient to cause broad-scale changes of ecosystems.
Precipitation variability during the Holocene has lasting
impact on the terrestrial landscape where water is stored
with a long residence time, that is, as lakes and glaciers. The
effects of precipitation variability on soil water and vegetation are fleeting because there is little interannual water storage, especially in shrublands and forests whose growth represents the integration of climate over long time scales.
Holocene geomorphic evidences may more reflect shortterm climatic singularities than broad structural changes in
global circulation. Enzel et al. (1992) proposed that the concurrent snow accumulation in Sierra Nevada, high precipitation, stream flows in southern California, and high lake stands
in Mojave River at 1600 CE all can be explained from singular
flood-producing events during a single season. For example,
high snow accumulations in a few wet years in the high Sierra
Nevada will ultimately produce more uniform runoff into
Mono and Walker Lakes. A striking example is a large flood in
1605 based on the thickness of sea floor varves in the Santa
Barbara Channel (Schimmelmann et al. 1998) that is rivaled
only by two other events in the past 1,000 years, floods in
1414 and 1969. (In the latter, as much as 70–120 cm of precipitation fell in the Transverse Ranges in 9 days.) The 1605
flood is apparently reflected in other proxy data. Graumlich
(1993) concluded from tree-rings that 1602 was the end of a
1566–1602 drought. Scuderi (1990) described a sharp decline
in ring-width at this time. Using tree-ring width data for 1366
to 1985 from Pseudotsuga macrocarpa in the mountains in
coastal central California, Hasten and Michaelsen (1994)
found that 1604 was the fourth wettest year; 1601–1611 was
the third wettest 11-year period. But tree rings describe
changes in growth rates, not changes in vegetation.
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Areas for Future Research
Askin, R.A. 1992. Late Cretaceous-early Tertiary Antarctic outcrop
evidence for past vegetation and climates. In: The Antaractic
The future of paleobotany and paleoecology bodes well with
the emergence of stable isotope geochemistry which has
made possible the independence of data on fossil floras and
proxy data on climate. The marine stable isotope record now
provides a well constrained time line over the period of evolution of modern assemblages from greenhouse climates in
late Cretaceous-early Tertiary to modern ice house climates.
The research agenda should continue in similar directions.
Studies of new and existing macrofossil and microfossil
assemblages can be assessed on their own terms with respect
to plant evolution and migration (Sauer 1988), and can be
correlated with emerging climatic proxy evidences from geochemistry, seen in a variety of such sediments as lake beds,
marine shell deposits, varved sediments, wood of ancient
trees, as well as physical evidences in lake stands, glaciation,
stabilized sand dunes, and fossil charcoal left by fires. Perhaps most frustrating is the dearth of macrofossil data from
packrat middens in the mediterranean climates of California, due possibly to poor preservation in mesic climates
compared to the rich record in deserts. Global climate models (GCMs) have elegant physics which provide scenarios of
past global atmospheric and oceanic circulations, but models are always incomplete. Hence GCMs will be bolstered by
empirical guidance. While global climate models help constrain climatic inference from proxy data, conflicting proxy
data will encourage the investigation of model assumptions.
Answers concerning the effects of recent global warming are
unlikely to come from deductive global climate models, but
rather from the oceans which are the primary heat reservoirs
of the planet. Increasing the empirical domain from the
kinds of evidences given above will slowly but surely lead to
important advances in the paleobotany and paleoecology in
California, as well as our understanding of the modern flora
and vegetation.
Environment: Perspective on Global Change. Antarctic Research
Series 56: 61–73.
Axelrod, D. I. 1937. A Pliocene flora from the Mount Eden beds,
southern California. Carnegie Inst. Wash. Contrib. Paleotol. 476:
125–183.
Axelrod, D.I. 1950. Further studies of the Mount Eden flora, southern
California. Carnegie Inst. Wash. Contrib. Paleontol 590: 73–117.
Axelrod, D.I. 1966. The Pleistocene Soboba flora of southern California. University of California Publications in Geological Sciences 60: 1–79.
Axelrod, D.I. 1973. History of the Mediteranean ecosystem in California. ed. F. DeCastri and H. A. Mooney, Mediterranean type
ecosystems: Origin and structure. Springer-Verlag.
Axelrod, D.I. 1980. History of the maritime closed-cone pines, Alta
and Baja California. University of California Publications in Geological Sciences Vol. 120. University of California Press. Berkeley,
CA.
Axelrod, D.I. 1981. Holocene climatic changes in relation to vegetation disjunction and speciation. American Naturalist 117: 847–870.
Axelrod, D. I. 1988. Outline history of California Vegetation. Pp.
139–193 in: Terrestrial Vegetation of California. ed. M.G. Barbour
and J. Major. California Botanical Society. Davis, CA.
Bains, S., R.M. Corfield, and R.D. Norris. 1999. Mechanisms of climate warming at the end of the Paleocene. Science 285: 724–727.
Barbour, M.G., and R.A. Minnich. 1990. The myth of chaparral convergence. Israel Journal of Botany 39: 453–463.
Barbour, M.G., N.H. Berg, T.G.F. Kittel, et al. 1991. Snowpack and
the distribution of a major vegetation ecotone in the Sierra
Nevada of California. Journal of Biogeography 18: 141–149.
Barnett, T.P, D.W. Pierce, M. Latif, et al. 1999. Interdecadal interactions between the tropics and mid-latitudes in the Pacific Basin.
Geophysical Research Letters 26: 615–618.
Bartlein, P.J., K.H. Anderson, P.M. Anderson, et al. 1998. Paleoclimate
simulations for North America over the past 21,000 years; features
of the simulated climate and comparisons with paleoenvironmental data. In: Late Quaternary climates; data synthesis and model
experiments (T. Webb III, ed.). Quaternary Science Reviews 17:
549–585.
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