(This is a sample cover image for this issue. The actual cover is not yet available at this time.)
This article appeared in a journal published by Elsevier. The attached
copy is furnished to the author for internal non-commercial research
and education use, including for instruction at the authors institution
and sharing with colleagues.
Other uses, including reproduction and distribution, or selling or
licensing copies, or posting to personal, institutional or third party
websites are prohibited.
In most cases authors are permitted to post their version of the
article (e.g. in Word or Tex form) to their personal website or
institutional repository. Authors requiring further information
regarding Elsevier’s archiving and manuscript policies are
encouraged to visit:
http://www.elsevier.com/copyright
Author's personal copy
Quaternary Science Reviews 40 (2012) 89e106
Contents lists available at SciVerse ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
An organic geochemical record of Sierra Nevada climate since the LGM from
Swamp Lake, Yosemite
Joseph H. Street a, b, *, R. Scott Anderson c, Adina Paytan b
a
Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94305, USA
Institute of Marine Sciences, University of California, Santa Cruz, CA, USA
c
Center for Environmental Sciences and Education and Quaternary Sciences Program, Northern Arizona University, Flagstaff, AZ, USA
b
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 23 April 2011
Received in revised form
27 January 2012
Accepted 28 February 2012
Available online xxx
Sediment records from Swamp Lake (SL) in the central Sierra Nevada, California, provide evidence of
climatic change on millennial and centennial timescales over the last w20,000 years. Total organic carbon
(TOC) abundance varied in concert with elemental and isotopic tracers of organic matter (C/N, d13Corg, d15N),
biogenic silica content, total magnetic susceptibility, and sediment lithology. We interpret the down-core
proxy records as representing the response of the lake environment, in terms of temperature, seasonal ice
cover, mixing regimes, runoff and in situ OM and nutrient cycling, to shifting climate states. These environmental factors in turn drove changes in algal productivity, OM sources, microbial OM regeneration and
secondary production, and detrital input. The late Pleistocene (w19.7e10.8 cal. kyr BP) was dominated by
fluctuations between relatively warm/dry intervals with high TOC (17.4e16.5, 15.8e15.0, 13.9e13.2,
11.4e11.0 cal. kyr BP) and cold/wet intervals (16.5e15.8, 14.8e13.9, 13.1e11.6, 11.0e10.7 cal. kyr BP) characterized by low TOC and high detrital input. The Holocene (w10.7 cal. kyr BP e present) was characterized
by three abrupt increases in TOC (after w10.8, 8.0, and 3.0 cal. kyr BP) and numerous century-scale fluctuations. TOC increases reflected enhanced lake productivity and OM recycling, and reduced detrital input,
in response to changing winter temperature and hydrologic regimes. Inferred environmental changes at SL
correlate with other Sierra Nevada paleorecords, and with reconstructed sea surface temperatures along
the California margin. Parallel changes in the SL and SST records over the past w20,000 years provide new
evidence that continental climate in the Sierra Nevada and the California Current system have responded,
on multiple timescales, to common drivers in North Pacific ocean-atmospheric circulation.
Ó 2012 Elsevier Ltd. All rights reserved.
Keywords:
Sierra Nevada, California
Paleoclimate
Paleolimnology
Sedimentary organic matter
Stable isotopes
1. Introduction
The pacing of natural climatic variability is of particular
importance in California, where a limited water supply is dependent on snowmelt runoff from the Sierra Nevada, and is thus
sensitive to changes in mean values and seasonal patterns of
precipitation and temperature (Wilkinson et al., 2002; Miller et al.,
2003). In light of the short instrumental record, only paleoclimatic
records can reveal patterns of variability operating on timescales
longer than a few decades. These patterns may influence regional
water supplies and the occurrence of extreme events such as floods
* Corresponding author. Department of Geological and Environmental Sciences,
Stanford University, Stanford, CA 94305, USA. Tel.: þ1 415 298 2543; fax: þ1 650
725 2199.
E-mail address: [email protected] (J.H. Street).
0277-3791/$ e see front matter Ó 2012 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2012.02.017
and droughts, and obscure or intensify the effects of human interference with global climate.
Climate variability in California and western North America is
affected by Pacific Ocean surface conditions that influence basinwide atmospheric circulation patterns (Seager and Vecchi, 2010).
For example, interannual cycles in tropical Pacific sea surface
temperature (SST) (El Niño-Southern Oscillation, ENSO) are associated with anomalous temperature and precipitation throughout
western North America and warm SST off California (e.g., Redmond
and Koch, 1991; McGowan et al., 1998; Dettinger et al., 2001; Diaz
et al., 2001; Higgins et al., 2002; Castello and Shelton, 2004).
Likewise, decadal shifts in North Pacific SST and atmospheric
circulation are linked to changes in regional temperature, precipitation, snow accumulation, streamflow and runoff timing, and
related ecological processes (Dettinger and Cayan, 1995; Mantua
et al., 1997; McCabe and Dettinger, 2002; Millar et al., 2004;
Tootle et al., 2005; Hunter et al., 2006). The regional influences of
Author's personal copy
90
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
ENSO and the Pacific Decadal Oscillation (PDO) have been detected
in paleorecords spanning the last millennium (e.g., Biondi et al.,
2001; Benson et al., 2003; Cook et al., 2004; MacDonald and
Case, 2005; Graham et al., 2007; Herweijer et al., 2007; Conroy
et al., 2009), but our understanding of the relationship between
ocean conditions and climate in western North America over longer
timescales, and during climate regimes of the more distant past,
remains fragmentary (Barron and Anderson, 2011). This is particularly true of the large Sacramento-San Joaquin River watershed
draining the Sierra Nevada.
To date, only a handful of published paleoclimatic records from
sites in the Sierra Nevada region combine high temporal resolution
(centennial-scale and finer) and length (full Holocene or longer)
(e.g., Hughes and Graumlich, 1996; Benson et al., 2002; Mensing
et al., 2004). Previous research has linked millennial-scale
climatic and environmental change to insolation, continental ice
sheet dynamics and local glacial advances/retreats, and various
ocean processes (e.g., Davis et al., 1985; Davis and Moratto, 1988;
Anderson, 1990; Smith and Anderson, 1992; Anderson and Smith,
1994; Koehler and Anderson, 1994, 1995; Benson et al., 1996,
1997, 1998; Davis, 1999a, b; Mensing, 2001; Brunelle and Anderson,
2003; Porinchu et al., 2003; Mensing et al., 2004; Bacon et al., 2006;
Negrini et al., 2006; Oster et al., 2009). However, these studies have
only occasionally addressed variability on sub-millennial timescales during climate regimes prior to the late Holocene.
In the present study we use organic geochemical analyses of
bulk sediments from Swamp Lake (SL), a mid-elevation lake in
Yosemite National Park (Fig. 1), to reconstruct paleoenvironmental
variability in the Sierra Nevada at w100-year resolution over the
past w20,000 years. Lake sedimentary organic matter (OM) derives
from both autochthonous production (algae, bacteria and aquatic
plants) and the input of terrestrial plant material from the
surrounding watershed, and is thus an important archive of paleoenvironmental change in the lake basin. Carbon and nitrogen
elemental abundances (TOC, TN, C/N) and isotopic compositions
(d13Corg, d15N) in sedimentary OM respond to changes in the
balance of OM sources, vegetation distribution, primary production
and lake trophic status, nutrient inputs and sources, and carbon
utilization and cycling (e.g., Meyers and Takemura, 1997; Hodell
and Schelske, 1998; Brenner et al., 1999; Meyers, 2003). These
proxies may therefore provide sensitive indicators of climatedriven environmental change (Meyers and Takemura, 1997;
Meyers and Lallier-Verges, 1999; Talbot and Laerdal, 2000). At SL,
we apply OM proxies to gather new insight into the Late
PleistoceneeHolocene climate history of the Sierra Nevada. In
particular, we attempt to better define the major climatic regimes
of the last 20 kyr, patterns of century-scale variability within these
regimes, and their relationship to existing reconstructions of SST in
the eastern North Pacific (e.g., Seki et al., 2002; Barron et al., 2003).
The SL record provides a western Sierra point of comparison with
records from eastern Sierra drainages (Benson et al., 1996, 1997,
1998, 2002, 2010; Davis, 1999a; Mensing et al., 2004; Briggs et al.,
2005; Bacon et al., 2006), and contributes to the synthesis of
a more complete picture of regional paleoclimate in western North
America since the last glacial maximum (LGM).
2. Materials and methods
2.1. Study area
Swamp Lake is a mid-elevation tarn lake located in Yosemite NP
in the central Sierra Nevada (1554 m elevation, 37 570 N, 119 490 W)
(Fig. 1). Thanks to its location less than 2 km east of the maximum
westward (and downslope) edge of Tioga stage glaciation (Smith
and Anderson, 1992), SL was ice-free earlier than high-elevation
sites, and the sediments accumulating here provide the longest
continuous paleoenvironmental record yet recovered from the
western Sierra. The granitic SL basin covers w1.3 km2, while the
lake itself has a surface area of 0.08 km2 and reaches a maximum
depth of 20 m. Inlet and outlet streams to the lake are ephemeral,
operating only from the winter-wet season to early summer,
though perennial subsurface inflow and outflow may occur. At
present, SL does not freeze during the winter, with water temperatures reaching a minimum of w5 C (Fig. 2a). Continuous air
temperature readings taken in DeceFeb of 2006e2007 fluctuated
between 7 and 5 C (Roach, 2010), with no below-zero excursion
Fig. 1. (a) Map of California, with locations of Swamp Lake (SL, star), Yosemite (green shading) and the sites of other paleorecords in the region: LT, Lake Tahoe; MC, Moaning
Cavern; ML, Mono Lake; OL, Owens Lake; PL, Pyramid Lake; TL, Tulare Lake. (b) Topographic map of the Swamp Lake drainage basin in northwestern Yosemite NP. Adapted from
Roach (2010). (c) Bathymetric map of Swamp Lake and coring locations. Contours in meters. (For interpretation of the references to color in this figure legend, the reader is referred
to the web version of this article.)
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
91
Fig. 2. Depth profiles of (a) temperature, (b) chlorophyll a, and (c) d13CPOM in Swamp Lake, 2008e2010.
and 4.1 C, respectively (1955e2008). Annual precipitation averaged 122 41 cm, with w90% occurring between November and
May. Annual snowfall averaged 293 109 cm (WRCC, 2012), representing perhaps 25e40% of total annual precipitation.
lasting long enough to induce lake freezing. Temperature stratification in SL develops over the course of the springesummer
growing season (Fig. 2a) following deep mixing/isothermal conditions during the winter. Chlorophyll a (chl a) profiles taken
between March and August (2008e2010) reveal a deep (10e15 m)
chl a maximum lasting through the summer, following an early
spring (MareApr) bloom in the epilimnion (Fig. 2b).
Terrestrial vegetation at the site is of the Sierra lower montane
forest type. Common species include ponderosa pine (Pinus ponderosa), sugar pine (Pinus lambertiana), incense cedar (Calocedrus
decurrens), white fir (Abies concolor), manzanita (Arctostapylos spp.)
and black oak (Quercus kellogii). Riparian species along the lake
margin include western azalea (Rhododendron occidentale),
mountain alder (Alnus rhombifolia), willow (Salix spp.) and deerbrush (Ceanothus integerrimus), while the western marsh and
shallow waters around the lake support an array of emergent,
floating, and submerged aquatic plants. Common emergent species
include three-way sedge (Dulichium arundinaceum), bogbean
(Menyanthes trifoliata), tule (Schoenoplectus acutus) and Carex
sedges; common aquatic plants include watershield (Brasenia
schreberi), floating pondweed (Potamogeton natans), yellow pond
lily (Nuphar lutea), and swaying bulrush (Schoenoplectus
subterminalis).
The Mediterranean climate of the western Sierra Nevada is
defined by warm, dry summers and cool, wet winters, and is governed by seasonal atmospheric circulation patterns in the North
Pacific. During the winter months, the Aleutian Low strengthens
and the Pacific Subtropical High moves equatorward, displacing the
westerlies over northern California and allowing storms originating
in the North Pacific to reach the Sierra Nevada. During the summer,
a strengthened, northward-trending Subtropical High blocks
moisture-bearing Pacific air masses from reaching California. At
Cherry Lake (w7 km NW of SL, w100 m lower in elevation),
summer (JJA) and winter (DJF) temperatures have averaged 20.4 C
2.2. Sediment cores
A series of sediment cores was recovered from Swamp Lake in
June 2002, as described by Anderson (2011). In this study we used
a 9-drive, 947 cm Livingstone core (Core 5) and a frozen finger (Core
8a) of the upper 83 cm, collected at a depth of 19.7 m (Fig. 1c). Cores
5 and 8a were composited based on the uppermost radiocarbon
ages in each core (Table 1) and a co-occurring tephra (57 cm in Core
8a) (Starratt and Anderson, in press). Sampling occurred at w5 cm
intervals throughout the core, but more frequently (1e3 cm) over
the upper 130 cm. A single sample was also collected from the topmost sediments of another freeze core (06-05, Roach, 2010), which
recovered sediments from the last w30 years that were missing
from Core 8a (Anderson, 2011). Cores 8a and 06-05 were correlated
based on measured 137Cs peaks (1964 C.E.) and several charcoal
layers traceable to 20th century fires in the vicinity of SL.
2.2.1. Age control
Age control for this study is provided by thirteen AMS radiocarbon dates on bulk sediment OM (Table 1), a tephra geochemically identified as the Mazama ash at 400 cm (7.6 kyr BP) (Hallett
et al., 1997; Anderson, 2011), and in the upper sediments of the
freeze cores, measurements of 137Cs (peak at 1964 C.E.) and
a charcoal layer representing a large local fire in 1996, verifiable by
counting of visible varves (Roach, 2010; Anderson, 2011). Radiocarbon ages were converted to calibrated ages before 1950 C.E.
(cal. yrs BP) using CALIB 5.0.2 and the IntCal04 calibration dataset
(Reimer et al., 2004; Stuvier et al. 2005). An ageedepth model was
developed using linear interpolation between calendar ages
Table 1
AMS radiocarbon age data for Swamp Lake cores 02-5 and 02-8a. All ages measured on bulk sediment.
Lab ID
Core
Composite core depth (cm)
14
WW4885L
WW4887L
WW4888L
WW4889L
WW4890L
WW4891L
WW4892L
WW4893L
WW4894L
WW4895L
WW4884L
Beta 202661
WW4886L
02-5
02-5
02-5
02-5
02-5
02-5
02-5
02-5
02-5
02-5
02-5
02-5
02-8a
63.5
79.5
123.5
169.7
221.8
250.9
566.6
607.6
619.7
706.5
717
865.5
60
566
969
1912
2683
3196
3871
9495
9929
10,479
12,152
12,640
14,810
585
C age (yr BP)
42
42
35
37
38
46
52
53
58
63
40
70
44
Calendar age (cal. yr BP)
Error (2s) (cal. yr)
595
862
1859
2789
3418
4303
10,775
11,343
12,477
14,007
14,917
18,006
600
36
84
81
52
65
101
179
137
188
186
304
219
54
Author's personal copy
92
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
(radiometric, tephra, charcoal) (Fig. 3). As shown in Table 1, 2s
errors for the calibrated ages range from 36 to 304 years, with
greater age uncertainty in the lower part of the core. All ages given
in the text refer to calibrated, calendar ages. Below the oldest
radiocarbon date at 865 cm the ageedepth relationship is unconstrained; dates from this interval (865e947 cm) should be treated
with caution. The small size of the SL catchment and the absence of
carbonates in the bedrock and lake sediments minimize the risk of
age errors associated with the use of bulk sediment for radiocarbon
analysis.
2.3. Field sampling
To determine the C and N isotopic values and C/N ratios of
potential sources of OM to SL sediments, samples of the most
common terrestrial and aquatic plant species at SL were collected
during the growing season (MayeOct., 2007e2010). Suspended
particulate OM (POM) was also collected at multiple depths on precombusted Whatman GF/F filters. Plant and POM samples were
stored on ice in the field and frozen upon returning to the lab.
Depth profiles of lake water temperature were taken in the central
basin using a YSI 85/50 hand-held probe (Fig. 2a). Water samples
(100e300 mL) for chl a analysis were collected along depth transects (Fig. 2b), filtered through GF/F filters and extracted in vials
containing 90% acetone. The vials were stored on ice in the dark in
the field and placed in a 4 C refrigerator after returning to the lab.
After 24 h, the samples were analyzed on a Turner fluorometer
before and after acidification for phaeopigment correction.
2.4. Sample analysis
Sediment, plant, and POM samples were dried, homogenized
and weighed into tin capsules for analysis. Total organic carbon
(TOC), total nitrogen (TN), d13Corg, and d15N were determined using
a PDZ Europa ANCA-GSL elemental analyzer interfaced to a PDZ
Europa 20-20 continuous flow isotope ratio mass spectrometer
(Sercon Ltd., Cheshire, UK) in the dual-isotope mode. Carbon and
nitrogen isotope ratios are expressed in conventional delta (d)
notation as the per mil deviation (&) from the Vienna PeeDee
Belemnite (VPDB) carbonate rock standard and from atmospheric
N2, respectively. Instrumental precision, based on laboratory standards and measured as the standard deviation of all standard
analyses (n ¼ 115) was 0.05& for d13Corg and 0.2& for d15N. The
reproducibility of TOC and TN measurements, estimated as the
pooled standard deviation of duplicate analyses (n ¼ 66, 0.1e21 wt
% C), was 6.8% (avg. 0.4 wt% C) and 9.3% (avg. 0.04 wt% N),
respectively. Though carbonates were not expected to be present in
SL sediments, a subset of sediment samples were fumigated in
a chamber with concentrated HCl (Harris et al., 2001) and
compared with unacidified duplicates. TOC and d13C values were
unaffected by acidification, which we took as evidence that
carbonates were absent or of negligible abundance in these sediments. Biogenic silica (BSi) was extracted from sediments with 10%
Na2CO3 and concentrations determined with a spectrophotometer
(Hach DR/2000) with an estimated precision of 9 mg/g (6%),
following Mortlock and Froelich (1989).
3. Results
3.1. Geochemical and isotopic signatures of vegetation & POM
C/N and d13C measurements of vegetation and POM samples
collected in the summers of 2007e2010 allow for the mapping of
distinct chemical signatures for potential OM sources in the
modern lake basin (Fig. 4). Terrestrial vegetation was characterized
by high C/N ratios (40.1 24.6 (1s), n ¼ 15) and intermediate
d13Corg values (28.8 1.6&), submerged/floating aquatic vegetation had intermediate C/N ratios (16.4 2.9, n ¼ 8) and higher
d13Corg values (24.9 1.9&), with emergent aquatic plants plotting in between (C/N ¼ 25.2 10.2; d13Corg ¼ 26.9 1.0&; n ¼ 7).
Lake POM had low C/N (9.7 2.9, n ¼ 29) and d13Corg
(34.6 2.8&) values. The C/N and d13Corg values of these OM
source groups were statistically distinguishable from one another
above the 99% confidence level (p 0.01, ManneWhitney U-test)
with the exception of emergent and terrestrial plants. POM samples
from the hypolimnion (>6 m) had very low d13C values
(37.2 1.9&; n ¼ 10), distinct from epilimnion (32.8 1.9&,
n ¼ 12) and thermocline (33.2 2.0&, n ¼ 6) samples (Fig. 4). The
d15N values of potential OM sources were less unique and more
variable than their C/N and d13C values, with no clear distinction
between lake POM and vascular plants. Terrestrial, emergent, and
submerged/floating aquatic plants had d15N values of 2.6 2&,
0 1.3&, and 3.2 1.8&, respectively. The d15N of lake POM
(0.2 3.6&) ranged from 12.1 to þ4.0&.
3.2. Sediment lithology
Fig. 3. Ageedepth model for SL composite core.
The SL composite core is dominated by fine sediment units,
distinguishable from one another by OM content (low OM clays/
silts vs. high OM gyttjas) and structure (laminated vs. massive)
(Figs. 5and 6a). The late Pleistocene portion of the record
(947e560 cm; 19.7e10.7 kyr BP) consists largely of alternating beds
of clay/silt and gyttja (massive or laminated), with multiple sand
lenses and tephras. More significant sand layers occur at
658e642 cm (13.2e12.9 kyr BP) and 581e572 cm
(11.0e10.8 kyr BP). A high OM peat layer occurs at 624e616 cm
(12.6e12.1 kyr BP). The Holocene section of the core (559e0 cm;
10.7e0 kyr BP) consists of laminated gyttja, with occasional
narrow bands of massive gyttja, clay, or sand. A peat layer occurs
from 411 to 401 cm (7.8e7.6 kyr BP). A thicker interval of massive
gyttja occurs in sediments sampled by the freeze core (48e19 cm;
425e112 yr BP). Thin-section analysis indicates that the
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
93
Cal. yr BP
(± 2σ error)
595 ± 36
862 ± 84
1859 ± 81
2789 ± 52
3418 ± 65
4303 ± 101
7597 ± 54
(Mazama Ash)
10775 ± 179
11343 ± 137
12477 ± 188
14007 ± 186
14917 ± 304
Fig. 4. d13Corg vs. C/N of modern OM sources (large solid-color symbols:
plants ¼ squares, POM ¼ circles) and down-core sediments (small symbols). Average
values and first standard deviation (error bars) shown for modern end-members. (For
interpretation of the references to color in this figure legend, the reader is referred to
the web version of this article.)
laminations present throughout the SL cores represent annual
varves (Roach, 2010; Anderson, 2011) consisting of lightedark
couplets. Light layers are dominated by diatom frustules deposited during warm-season blooms, while dark layers are a mixture of
OM, charcoal and clastic fragments associated with wet-season
runoff into the lake (Roach, 2010).
13
15
3.3. Down-core variations in TOC, TN, C/N, d Corg, d N,
BSi and lithology
Large changes in the amount and composition of OM accumulating in SL sediments have occurred over the past 20,000 years,
signaling large rearrangements of the surrounding environment
(Fig. 6bef). Organic content (TOC, TN) increased 100-fold from the
base of the core (0.27 wt% C, 0.1 wt% N) to the present (27 wt% C,
2.8 wt% N). Patterns of TOC and TN accumulation were nearly
identical throughout the core (r ¼ 0.97, p < 0.0001), and henceforth
we generally discuss only changes in TOC. Superimposed on the
long-term trend was a great deal of variability on multiple timescales, which is also reflected in the C/N ratio (6.9e21.1), d13Corg
(22.2& to 35.5&), d15N (2.8& to 2.0&) and BSi content
(5e556 mg/g) of sediments. The long-term average C/N ratio
increased between 19.7 and 15.0 kyr BP, and declined toward the
present. Long-term average d15N values declined between 19.7 and
14.0 kyr BP, increased from 14.0 to 10.8 kyr BP, and declined
through most of the Holocene.
Fluctuations in TOC, C/N, d13Corg, and d15N were significantly
associated with one another (Fig. 4, Fig. 7aee; Table 2) through
most of the SL record. TOC was negatively related to C/N, d13Corg,
18006 ± 219
laminated gyttja
massive gyttja
peat
gray clay/silt
layered clay
sand
tephra
Fig. 5. Generalized sediment stratigraphy for the SL composite core.
and d15N, which in turn were positively correlated (Table 2). Deviations from these relationships occurred during certain intervals,
most notably the peat at 411e401 cm (7.8e7.6 kyr BP), several
extreme low TOC intervals of the late Pleistocene (w19.7e18.2, 17.3,
14.5, 16.4e16.2,10.8 kyr BP), and the high TOC interval between 751
and 721 cm (15.8e15.0 kyr BP, SLZ-4). TOC covaried with BSi
(Fig. 6g) throughout the record, but were not compared quantitatively because the analyses were not done on the same samples.
Loss-on-ignition measurements of OM content (Anderson, 2011)
were positively related to BSi, (r ¼ 0.57, p < 0.001, not shown).
3.3.1. Late Pleistocene (19.7e10.7 kyr BP)
The Late Pleistocene (LP) section of the core (947e560 cm,
19.7e10.7 kyr BP) contains nine alternating low- and high-TOC
intervals (Fig. 6; “Swamp Lake Zones”, SLZ, 1e9) that generally
Author's personal copy
94
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
Fig. 6. Sediment proxy records from SL composite core over the last 19,700 years. (a) Sediment stratigraphy; (b) total organic carbon (TOC); (c) total nitrogen (TN); (d) biogenic silica
concentration (BSi); (e) organic carbon/total nitrogen molar ratio (C/N); (f) d13Corg; (g) d15N; (h) magnetic susceptibility. TOC maxima are highlighted in gray. Holocene intervals of
rapid TOC increase are indicated by cross-hatching. Holocene TOC minima (d13Corg maxima) are assigned letters “a”e“s”.
correspond to changes in sediment type. The earliest interval (SLZ1, 947e878 cm, 19.7e18.2 kyr BP) is composed of gray clay and silt
with low organic content (<1 wt% C, <0.12 wt% N), low C/N (7e12),
high d13Corg (23.3 to 28.5&) and d15N (0e1.4&) (Fig. 6bef).
Following SLZ-1 were eight alternating TOC maxima (centered at
w16.9, 15.3, 13.5, 11.1 kyr BP) and minima (w16.4, 14.5, 12.6, and
10.8 kyr BP).
All high TOC intervals consisted of massive and laminated gyttja
(Fig. 6a), but nonetheless differed in important respects. During the
earlier TOC maxima (SLZ-2b, 839e797 cm, 17.5e16.5 kyr BP; SLZ-4,
768e717 cm, 15.8e15.0 kyr BP), moderately high TOC (6e12 wt%)
corresponded to distinct C/N maxima (14.6e21.1) and d13Corg
minima (32.8 to 30&). The latter TOC maxima (SLZ-6,
702e662 cm, 13.9e13.2 kyr BP; SLZ-8, 608e587 cm,
11.4e11.0 kyr BP) reached higher peak values (15.3 and 16.8 wt%
C, respectively), but with C/N (14.7e20.9) and d13Corg values (32 to
25.6&) less distinct from neighboring low TOC intervals. The LP
high TOC intervals generally corresponded to local BSi maxima
(especially SLZ-4, 331e469 mg/g).
LP TOC minima (SLZ-3, 794e767 cm, 16.5e15.9 kyr BP; SLZ-5,
716e702 cm, 14.9e13.9 kyr BP; SLZ-7, 655e612 cm,
13.1e11.8 kyr BP; SLZ-9, 583e560 cm, 11.0e10.7 kyr BP) ranged
from 0.4 to 3.2 wt% C at their lowest values, and contained a more
varied array of sediment types than TOC maxima. SLZ-3 consisted of
massive and layered gray clay, similar to that of SLZ-1, while SLZ-5
(laminated gyttja alternating with clay and sand layers), SLZ-7
(sand, gyttja, peat) and SLZ-9 (massive þ laminated gyttja, interbedded sand) were more lithologically complex. SLZ-7 contained
several distinct zonations: Coarse sand and gravel (655e642 cm);
laminated gyttja (642e637.5 cm); massive gyttja with wood and
sand (637.5e624 cm); and peat (624e616 cm). SLZ-3 and SLZ-5
were characterized by locally low (high) C/N (d13Corg) values,
while SLZ-7 and SLZ-9 were associated with high or intermediate C/
N and d13Corg values. TOC and BSi minima often corresponded (e.g.,
at w16.3, 14.0, 12.8 and 10.9 kyr BP). Following SLZ-7 (after
w11.8 kyr BP), the dominant timescale of variability appears to
have decreased; SLZ-8 and -9 could alternatively be described as
a series of century-scale fluctuations leading into the early
Holocene.
3.3.2. Holocene (10.7 kyr BPepresent)
The Holocene section of the SL core consists of laminated gyttja
with brief intervals of sand, clay, or massive gyttja (Figs. 5and 6a).
Sedimentary OM content was typically higher than during the LP
(5e27 wt% C, 0.3e2.8 wt% N) but quite variable on centennial to
millennial timescales (Fig. 6b, c). TOC increased over the Holocene
in three relatively abrupt steps at 10.8e10.5, 8.0e7.4, and
3.0e2.8 kyr BP, corresponding to increased BSi and lowered C/N,
d13Corg and d15N (Fig. 6ceg). Based on these proxy shifts we defined
early (EH, 10.7e8.1 kyr BP, w12 wt% C), middle (MH, 7.5e3.1 kyr BP,
w16 wt% C) and late Holocene (LH, 3.0 kyr BPepresent, w17 wt% C)
sub-intervals, during which TOC fluctuated around relatively
stationary long-term means. TOC was strongly anti-correlated with
both C/N (11.2e19.4) and d13Corg (35.5& to 22.2&) throughout
the Holocene; C/N and d13Corg were positively correlated, and
typically lower than LP values. d15N decreased by >1.5& between
10.8 and 1.4 kyr BP, before increasing by a similar amount between
1.4 and 0.3 kyr BP. Higher frequency variability in d15N, however,
was muted in comparison to the other proxies.
EH sediments had the lowest TOC, (9e17 wt% C) and highest
baseline C/N (12.9e18.6), d13Corg (31 to 29&) and d15N (1.1 to
0.2&) values of the Holocene interval. Low BSi values (27e150 mg/
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
95
Fig. 7. Relationships among bulk OM proxies. (a) C/N v. TOC; (b) d13Corg v. TOC; (c) d15N v. TOC; (d) d15N v. C/N; (e) d15N v. d13Corg. Intervals represented by distinct symbols: SLZ-1,
19.7e18.2 kyr BP (C); “other low TOC, ( ); SLZ-4, 15.8e15.0 kyr BP (þ); SLZ-7, ( ); peat at 412e402 cm, 7.8e7.6 kyr BP (); tephra at 108e111 cm, w1.5 kyr BP ( ).
g, 10.7e9.8 kyr BP), increased (to w200 mg/g) after 9.5 kyr BP, in
parallel with declining d15N. Low TOC (low BSi, high C/N, d13Corg)
events occurred at 10.0, 9.4 and 8.9 kyr BP (Fig. 6b, events seq). The
EH ended abruptly with a w10 wt% decline in TOC and deposition
of a sand layer at w8.05 kyr BP, followed by deposition of high TOC
(w23 wt% C) peat (7.7e7.6 kyr BP). This EHeMH transition
(8.0e7.5 kyr BP) appears as a broad peak in the C/N, d13Corg and d15N
records.
The MH (7.5e3.1 kyr BP) at SL was expressed as an extended
period of high TOC (11.3e19.7 wt% C), on average 3e4 wt% greater
than during the EH. d13Corg (31.5 to 27.1&), C/N (12.1e16.5) and
d15N (1.4 to 0.3&) values were low relative to the EH, while BSi
content was high (150e382 mg/g). After w6.0 kyr BP, baseline C/
N values increased, while BSi and d15N values decreased. Higherfrequency variability was prominent, with eight parallel, multicentury fluctuations in TOC, C/N and d13Corg (Fig. 6). Notable TOC
minima (C/N, d13Corg maxima) occurred at 7.5, 7.2, 6.6e6.35, 5.7, 4.9,
4.4, 3.7, 3.25 and 3.0 kyr BP (Figs. 6 and 10, events nef). Aside from
the d15N maxima and BSi minima at w3.7 and 3.0 kyr BP, centuryTable 2
Correlations among SL OM proxies. All correlations are significant at the 99.9% level
(p < 0.0001).
TOC
Correlation coefficients (r)
TOC
1
C/N
e
13
d Corg
e
15
d N
e
Coefficients of determination (r2)
TOC
1
C/N
e
d13Corg
e
13
d N
e
a
C/N
d13Corg
d15N
0.63a
1
e
e
0.69a
0.70a
1
e
0.55a
0.51a
0.52
1
0.40
1
e
e
0.47
0.49
1
e
0.31
0.27
0.27
1
Peat, tephra, “low TOC” and/or SLZ-4 intervals excluded.
scale fluctuations in these records did not correspond closely to the
TOC, C/N and d13Corg records.
A distinct transition into the LH occurred between 3.0 and
2.8 kyr BP, consisting of large increases in TOC and BSi, and parallel
declines in d13Corg, C/N, and d15N. TOC remained high (10e27 wt
% OC) from 2.8 kyr BP to the present, but underwent three multicentury fluctuations (also seen in BSi, d13Corg and C/N), with TOC
maxima (d13Corg minima) at 2.7e1.8, 1.0e0.5 and 0 kyr BP. The most
dramatic LH changes occurred in the last 1000 years. From a broad
TOC maximum (d13Corg, C/N minimum) during the medieval period
(1000e550 yr BP), TOC declined (d13Corg, C/N increased) gradually
until the mid-15th century (w420 yr BP), after which an abrupt drop
in TOC (increase in C/N, d13Corg, d15N) coincided with deposition of
massive gyttja. During the last 150 years, TOC and BSi increased
three-fold, alongside large declines in C/N, d13Corg and d15N.
4. Discussion
4.1. Controls on sedimentary OM abundance
The OM abundance in SL sediments at any given time reflects
the balance of OM inputs (autochthonous & allocthonous) and
losses (heterotrophic respiration), as well as dilution by other
components, such as biogenic silica and inorganic detrital sediment. Each of these terms responds to climatic and environmental
changes. For instance, the productivity and composition of aquatic
organisms and terrestrial plant communities are sensitive to
temperature, growing season length, precipitation patterns, and
changes in lake level and mixing regimes. Likewise, the amount and
OM content of detrital input to the lake, and thus the extent to
which it dilutes (or enriches) sedimentary OM, will vary with
climate-driven changes in soil stability, vegetation cover, and runoff
in the drainage basin. The pathways and rates of OM degradation in
Author's personal copy
96
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
the water column and sediments are also strongly influenced by
environmental factors including productivity, temperature and
oxygen status. While the effects of detrital input on TOC can often
be inferred from sediment lithology and other proxies (e.g.,
magnetic susceptibility), it is not trivial to distinguish between the
effects of changing OM production, water column regeneration and
post-depositional diagenesis.
If OM losses related to post-depositional diagenesis dominated
the TOC record at SL, a monotonic down-core trend of decreasing
TOC might be expected. While TOC and TN both decrease drastically moving down-core (Fig. 6bec), this decrease was discontinuous, occurring abruptly at depths corresponding to times of major,
independently-established environmental changes (see x4.4e4.5).
Much of the variability in the proxy records occurs cyclically, which
is also difficult to explain in terms of diagenesis alone. In addition,
TOC and TN are strongly correlated throughout the record, ruling
out major selective losses (or retention) of one element over time.
The presence of varves through much of the SL record is indicative
of an anoxic depositional environment (Roach, 2010; Anderson,
2011), which would tend to increase OM burial efficiency and
preservation (Gong and Hollander, 1997; Lehmann et al., 2002;
Galman et al., 2008). Anaerobic respiration has been observed to
impart a distinct geochemical signature to residual OM, with
increased C/N and decreased d13Corg and d15N relative to primary
OM (from preferential degradation of N-rich, 13C and 15N-enriched
labile compounds) (Lehmann et al., 2002; Teranes and Bernasconi,
2005; Li et al., 2008). The strong positive correlations among C/N,
d13Corg and d15N in the varved portions of the SL record are thus
evidence that anaerobic diagenetic processes do not dominate the
proxy signals.
In light of these considerations, we conclude that the predominant controls on TOC abundance in SL sediments are (a) changes in
OM input, reflecting the productivity of lake algae, bacteria and
aquatic plants as well as allocthonous inputs from terrestrial plants,
(b) OM reworking prior to or soon after deposition, and (c) dilution
by inorganic components. Changes in OM sources and cycling
within the lake can be traced using C/N, d13Corg and d15N (x4.2),
while changes in sediment lithology help constrain detrital input.
Our interpretation of the bulk OM record is supplemented by
measurements of BSi concentration (Fig. 6g) and sediment
magnetic susceptibility (MS, Fig. 6h; Anderson, 2011). BSi is
commonly used as a proxy for the component of lake productivity
contributed by diatoms or other siliceous algae (e.g., Colman et al.,
1995; Hu et al., 2003). Patterns of TOC and BSi abundance are
correlated over the SL record, supporting our contention that lake
productivity is an important contributor to TOC, and indicating that
dilution by BSi is not the major control on TOC. In granitic Sierra
Nevada catchments, MS is dominated by the signal from magnetic
minerals in lightly-weathered clastic debris (Konrad and Clark,
1998; Benson et al., 2002). We thus interpret MS as a proxy for
detrital input and watershed stability, and as an indicator of dilution effects in the OM record. This interpretation is supported by
the observation that peak MS values are associated with low OM
clay, silt and sand intervals (Fig. 6).
4.2. Sources of OM to Swamp Lake sediments
4.2.1. Evidence from C/N ratios and d13Corg
Higher plants, lake algae and bacteria produce OM with distinct
C/N and d13C signatures, and these proxies and have been used
extensively to determine sources of sedimentary OM (e.g., Meyers,
1994; Prahl et al., 1994; Meyers and Takemura, 1997; Tenzer et al.,
1997; Talbot and Laerdal, 2000; Hollander and Smith, 2001). C/N
ratios of sedimentary OM primarily reflect the proportion of algal/
microbial (low C/N) and vascular plant (high C/N) derived material.
The original source signature can, however, be altered by the
preferential degradation of labile compounds during anoxic
decomposition (Teranes and Bernasconi, 2005; Li et al., 2008) or
masked by the presence of inorganic N in very low OM sediments
(Meyers, 2003). Differences in d13Corg arise from isotopic fractionations associated with different carbon fixation pathways (e.g., C3
vs. C4 plants), the 13C-content of inorganic carbon sources (e.g.,
atmospheric CO2, dissolved CO2, bicarbonate, methane), and
a variety of degradational processes (LaZerte and Szalados, 1982;
O’Leary, 1988; Hollander and Smith, 2001; Meyers, 2003; Li et al.,
2008). The d13C of lake organisms in particular is also sensitive to
other environmental variables, including productivity and growth
rates, water temperature, light availability, and long-term changes
in atmospheric CO2 concentrations (e.g., Hollander and McKenzie,
1991; Laws et al., 1995; Hodell and Schelske, 1998; Gu et al., 1999).
As shown in Fig. 4, the C/N and d13C values of modern OM
sources to SL sediments e including terrestrial and aquatic plants,
and lake POM e define a mixing space encompassing the C/N and
d13Corg values of most down-core sediment samples. Based on this
observation, we argue that sedimentary C/N and d13Corg are to
a large extent controlled by variable OM contributions from source
groups similar to those found at SL at present. Other controls may
have operated on these proxies, but the close correlations between
d13Corg and C/N throughout the record (Fig. 4) and the negative
relationship between d13Corg and TOC (Fig. 7b) suggest that the OM
source signal predominates.
Accordingly, sediment samples with low C/N ratios and low
d13Corg values contain a greater proportion of OM derived from lake
algae and/or bacteria, while high C/N and d13Corg indicate a greater
proportion of higher plant material. The main cluster of points in
Fig. 4 traces a continuum between the POM and aquatic (emergent
and submerged/floating) plant end-members, suggesting that OM
contributions from aquatic higher plants have generally exceeded
terrestrial plant inputs. Three clusters of points diverge from the
dominant trend: (a) Low C/N, high d13Corg values from SLZ-1 and
extreme low TOC intervals of SLZ-2, 3, 5 and 9, probably reflecting
algal production under C-limited, oligotrophic conditions
(x4.3.3.1.); (b) high d13Corg values from the peat layer at the EHeMH
transition, likely reflecting OM inputs from aquatic plants (x4.5.2);
and (c) high C/N, low d13Corg values from LP intervals (especially
SLZ-4), reflecting the greater relative importance of OM inputs from
terrestrial plants during these time periods (x4.4.1.1.).
4.2.2. Low d13Corg values suggest a distinct hypolimnion POM
source
The d13Corg values of modern POM (d13CPOM) (30.2 to 40.2&;
Figs. 2c and 4) and many down-core samples at SL (Fig. 4) are low in
comparison to the expected range for lacustrine algae utilizing
dissolved CO2 in equilibrium with the atmosphere (24 to 30&;
Meyers, 1994, 2003). Moreover, d13CPOM in the hypolimnion after
the onset of stratification is significantly lower than epilimnetic
d13CPOM (Figs. 2c and 4; x3.1). Low d13C in POM and surface sediments are observed in lakes where algal productivity is supported
to some degree by uptake of 13C-depleted, respired carbon e
produced by OM regeneration in the water column or methanogenesis (and subsequent methane oxidation) in anoxic sediments
or bottom waters e or where 13C-depleted secondary bacterial
biomass (e.g., methanotrophs) contributes substantially to the total
OM pool (Rau, 1978; Meyers and Ishiwatari, 1993; Hollander and
Smith, 2001; Bastviken et al., 2003; Teranes and Bernasconi,
2005; Li et al., 2008).
The presence of varved, unbioturbated sediments through much
of the SL record is indicative of persistent seasonal anoxia at the
sediment surface, consistent with field measurements of dysoxia
(DO <3 mg/L) below 10 m depth in the modern lake. Depth profiles
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
of temperature, chl a and d13CPOM at SL (Fig. 2) hint at a seasonal
progression during which summer stratification (Fig. 2a) and
productivity at depth (10 m; Fig. 2b) promote microbial respiration and oxygen depletion in the hypolimnion and surface sediments. Under such conditions, anaerobic respiration of OM
produces methane (d13C z 65 to 50&) (Whiticar, 1999;
Hollander and Smith, 2001), and thus a source of 13C-depleted
carbon for methane-oxidizing bacteria, which in turn produce 13Cdepleted CO2 that may support algal productivity in the hypolimnion (deep chl a maximum, Fig. 2b). The 13C-depleted DIC pool that
accumulates in the hypolimnion in the summer is transferred to the
epilimnion during winter mixing and leads to low d13C of POM
during the early spring bloom (Fig. 2c). The frequent occurrence of
methanogenesis at SL is also suggested by the presence of bacterial
hopanoid biomarkers with very low d13C values (90 to 40&)
throughout the sediment record (Street et al., unpubl. data). Low
d13C values in SL sediments therefore reflect (a) the effects of an
anoxic lake bottom on the DIC pool and subsequent primary
production, and (b) the direct input of 13C-depleted bacterial
biomass to the sediments.
In light of these considerations, we interpret the sedimentary
d13Corg record as a measure of the balance of allocthonous vs.
autochthonous (including both algal and bacterial production) OM
sources, as with C/N, but also as a gauge of deep-lake C cycling
intensity and oxygen status.
4.2.3. Evidence from d15N
Nitrogen isotope compositions (d15N) of sedimentary OM have
been used to infer OM and N sources in lakes (e.g., Hodell and
Schelske, 1998; Brenner et al., 1999; Talbot and Laerdal, 2000;
Herczeg et al., 2001), but application of this proxy is complicated by
the complex biogeochemistry of N in lake systems (Meyers, 2003).
In the SL record, positive correlations with d13Corg and C/N
(Fig. 7eef; Table 2) ight indicate OM-source control of sedimentary
d15N, but the distribution of d15N values in modern OM sources
(x3.1) would result in an interpretation contradicting that suggested by d13C and C/N (i.e., lower d15N indicative of larger OM
contribution from vascular plants). We suggest instead that the
initial d15N source signal has generally been altered by OM degradation and N cycling in the SL watershed and water column.
OM regeneration in the water column (x4.2.2) would be
expected to generate 15N-depleted ammonium (NHþ
4 ), which, if
incorporated into subsequent primary productivity, would yield
biomass with low d15N values (4 to þ1&; Meyers and Ishiwatari,
1993; Li et al., 2008). Secondary production by heterotrophic
bacteria in the water column or sediments may also contribute 15Ndepleted biomass to the sediments while altering the d15N signature of residual primary OM (Lehmann et al., 2002). Modern
d15NPOM averages w0&, but with a broad range and no clear spatial
or temporal pattern, perhaps reflecting a complex interplay
between N-utilization during primary production (increases
d15NDIN), remineralization (decreases d15NDIN) and secondary
microbial production. Low d15N values (3 to 0&) in most downcore sediments nonetheless support the idea that OM recycling
and microbial productivity influence the SL d15N record.
High d15N (>0&) in the SL record is often associated with high
C/N and d13Corg (and low TOC) (Fig. 6, Fig. 7eef), which we have
interpreted as reflecting a greater proportion of vascular plant OM
in the sediments. However, since high d15N values are not consistently observed in modern plant samples (7.0 to þ1.6&), we
suspect that high d15N in the sediment record instead reflects (a)
reduced generation of 15N-depleted DIN within the lake due to
weakened OM regeneration, and/or (b) the contribution of
partially-degraded, residual soil and marsh plant OM, which is 15Nenriched relative to fresh plant litter (Kendall, 1998). During the
97
SLZ-1 (and other extreme low TOC intervals), high d15N was associated with high d13Corg (Fig. 7f) but low C/N (Fig. 7e), which we
attribute to algal OM production under oligotrophic, N-limited
conditions (x4.3.3.1.).
4.3. Environmental interpretation of the SL record
4.3.1. Sediment changes & long-term development of the SL basin
Sedimentological changes in the SL record (x3.2) reflect the
changing influence of several factors, including landscape stability
and runoff, lake level, and autochthonous production of biogenic
material (OM, BSi). Clay deposition during SLZ-1 (Fig. 6) probably
records high meltwater runoff and detrital input from alpine
glaciers in the drainage (x4.3.2.1.). We attribute the shift to organic
sedimentation (gyttja) after w18.2 kyr BP (Fig. 6) to further glacial
retreat, increased landscape stability and reduced runoff/sediment
transport, and increased lake productivity in response to a warming
climate. The long-term decline in clastic sedimentation over the LP
(Fig. 5) probably reflects increasing vegetation density and reduced
erosion in the SL drainage, consistent with a trend toward higher
sedimentary C/N (to w15 kyr BP, Fig. 6e), and pollen evidence for
the gradual replacement of alpine steppe with conifer forest
(Anderson, 2011). Higher frequency changes in sedimentation,
discussed in detail in x4.4.1, reflect the interaction of changing
runoff and detrital input patterns with in situ biogenic production.
Within the gyttja intervals that dominate the SL record, knowledge
of the OM content, BSi and magnetic susceptibility are crucial for
inferring the relative importance of detrital vs. autochthonous
components, and thus formulating an environmental interpretation. Peat layers deposited between 12.5e12.2 and 7.8e7.6 kyr BP
may represent intervals of low lake level.
4.3.2. Controls on TOC at SL
Through most of the SL record, sedimentary TOC is significantly
correlated with other proxies (Fig. 7aec; Table 2), allowing us to
make inferences about the environmental conditions influencing
TOC abundance. With a few exceptions, TOC is negatively related to
C/N, d13Corg, and d15N, and positively related to BSi. High TOC is also
often associated with deposition of varved sediments (Fig. 6). We
conclude that high TOC abundance at SL is associated with some
combination of high lake productivity (and thus a high proportion
of OM from algal or microbial sources), low dilution by detrital
sediments, and lake bottom anoxia. Moreover, low d13Corg values
observed during many high TOC intervals (e.g., LH) suggest that OM
regeneration, methanogenesis and microbial productivity in the
hypolimnion and surface sediments are important factors
contributing to high sedimentary TOC. While microbial activity
may degrade a significant portion of primary OM during high TOC
intervals, it also must support a large contingent of secondary
biomass that is preserved in the sediments. The relative importance
of ecosystem productivity versus detrital dilution in determining
sedimentary TOC concentrations is difficult to disentangle, but
based on the more frequent occurrence of clastic sediments and
high MS values, dilution was probably a more significant factor
during the LP interval.
4.3.3. Environmental end-members
Lake responses to climate and environmental change are
mediated by multiple site-specific factors (e.g., size, depth, stratification and mixing regimes, ice cover, nutrient distribution, water
inputs) and are thus complex and variable in space and time
(Schindler, 1997). In the absence of long-term monitoring data from
SL, proposed links between climate forcing and past changes in
productivity and sedimentation remain speculative. Nonetheless,
we argue that by examining proxy values and relationships during
Author's personal copy
98
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
intervals of the SL record when external environmental conditions
are partially constrained, we can bolster our interpretive framework and reach a better understanding of the full record. Specifically, we contrast two periods, SLZ-1 (19.7e18.2 kyr BP) and the
20th century, as potential “end-member” states in a continuum of
environmental conditions at SL.
4.3.3.1. SL at the LGM. SLZ-1 (19.7e18.5 kyr BP), though poorly
constrained in time (x2.2.1), represents the early SL after the onset of
Tioga stage glacial retreat. Sedimentation consisted of low TOC, high
MS detrital clay and silt (Fig. 6), likely reflecting input from
a sparsely-vegetated post-glacial watershed. The combination of low
C/N, high d13Corg, and high d15N in SLZ-1 sediments (Figs. 6and 7) is
consistent with low productivity in an oligotrophic (possibly turbid)
lake with limited supplies of DIN and DIC, and minimal higher plant
contributions of OM. It is also possible that inclusion of inorganic N
has depressed the measured C/N ratio in these low OM sediments
(Meyers, 2003). LGM climate in the Sierra Nevada was colder than
present, particularly during winter, with significant downslope
shifts in vegetation assemblages (Koehler and Anderson, 1994, 1995;
Mock and Bartlein, 1995; Benson et al., 1998; Anderson, 2011). Cold
water temperatures, persistent seasonal ice cover, low rates of OM
regeneration and bacterial activity and low pCO2 (w190 ppm vs.
280 ppm pre-industrial; Barnola et al., 2003) would have exacerbated C limitation and contributed to high d13Corg values in sedimentary OM (Hollander and McKenzie, 1991; Laws et al., 1995; Gu
et al., 1999; Li et al., 2008). The sedimentary characteristics of SLZ1 may thus provide a “signature” for cold-climate, oligotrophic,
and high runoff/detrital sediment-dominated conditions at SL.
4.3.3.2. SL in the 20th century. Proxy values in 20th century sediments from SL reflect a lake environment very different from that of
the LGM (Fig. 6aeh). High concentrations of TOC (14e27 wt.%, Fig.
8a), TN (1.4e2.8 wt.%) and BSi (22e42 wt.%), along with low d13Corg
(31.3 to 35.5&), d15N (1 to 2.8&) and C/N (11e13) in varved
sediments are indicative of high rates of algal productivity, low
detrital input, predominantly autochthonous OM sources, sediment
anoxia, and vigorous OM regeneration and secondary microbial
production. We also conclude, based on field measurements, that
these ecosystem characteristics are associated with a monomictic,
seasonally-stratified mixing regime (Fig. 2a), absent or intermittent
seasonal ice cover (Roach, 2010), hypolimnetic oxygen depletion,
and deep (10 m) summer chl a and POM maxima (Fig. 2b).
The large changes observed in the proxy records during the last
150 years (Figs. 6 and 8) may also reflect major recent changes in
the lake environment. Since 1850, TOC, TN and BSi in SL sediments
have increased two- to four-fold, while C/N, d13Corg and d15N
decreased markedly (by 4, 8&, and 2.7&, respectively), with the
greatest changes generally occurring since w1950. While it is
conceivable that these trends are an artifact of OM diagenesis, the
directions of change in the individual proxies, as well as the
parallels between organic and inorganic (BSi) proxies, are better
explained by changes in the lake environment (i.e., increased
productivity, OM cycling).
Regional climate in the Sierra Nevada over the last w100 years
has been warm relative to the preceding Little Ice Age (LIA,
w1500e1850 CE; Graumlich, 1993; Clark and Gillespie, 1997; Lloyd
and Graumlich, 1997; Konrad and Clark, 1998), and has experienced both summer and winter warming since 1950 in response to
shifts in North Pacific atmospheric circulation and, possibly, global
climate change (Dettinger and Cayan, 1995; Cayan et al., 2001;
Higgins et al., 2002; Millar et al., 2004). Plotted in Fig. 8b is a time
series of composited JaneMar. temperature for seven central California stations (Table 3; updated/expanded from Dettinger and
Cayan, 1995), showing a warming of w2 C since 1950 that
Fig. 8. Comparison of (a) Swamp Lake sedimentary TOC during the 20th century with:
(b) JaneMar mean temperature anomalies composited from seven Sierra region
climate stations (Table 1). For each station, anomalies were calculated as the deviation
from the average Jan.eMar. temperature over the period of record. Only the Sacramento station was active prior to 1930. The solid black line represents the 9-yr running
mean for the composite dataset; gray lines represent 9-yr running means for the Hetch
Hetchy and Cherry Valley stations nearest Swamp Lake; and (c) Water-year (OcteSep)
precipitation % of average (1930e2010) for the San Joaquin Basin 5-Station Index (Cal
DWR, http://cdec.water.ca.gov/). We extended this record to WY 1907 using a smaller
number of available stations (1913e1929, 4 stations; 1911e12, 3 stations; 1907e1910, 2
stations). Solid line represents 9-yr running mean.
parallels the period of most drastic change in TOC at SL (Fig. 7a).
Over the same period, winter precipitation shows no trend
(Fig. 8c). The recent warming trend in the Sierra Nevada has
occurred contemporaneously with increased anthropogenic
atmospheric deposition of nitrogen and phosphorus, with potential fertilizing effects even in remote, undisturbed wilderness lake
basins like SL (Jassby et al., 1994; Clow et al., 2003; Sickman et al.,
2003). As discussed below, the combined effects of climate
Table 3
Weather stations used to define 20th century trends in winter temperature (Fig. 8b).
Station
NCDC Coop
station #
Latitude
( N)
Longitude
( W)
Altitude
(masl)
Period
of record
Sacramento
Nevada City
Tahoe City
Twin Lakes
Hetch Hetchy
Cherry Valley
Modesto
047633
046136
048758
049105
043939
041697
045738
38 350
39 150
39 100
38 420
37 570
37 580
37 390
121 300
121 010
120 080
120 030
119 470
119 550
121 000
6
847
1899
2438
1180
1451
27
1878e2006
1931e2006
1931e2006
1949e1999
1931e2006
1956e2006
1931e2006
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
warming and nutrient enrichment provide a plausible explanation
for recent changes in the proxy records and by extension, other
high TOC intervals in the SL record.
The connections between climate and OM abundance at SL are
most likely indirect, involving feedbacks among lake mixing
regimes, winter ice cover, and nutrient cycling in the lake system.
Reduced winter ice cover, for example, has been observed to
enhance total productivity in lakes by influencing the timing of lake
overturning, the onset of stratification, and the length of the
growing season (Strub et al., 1985; Goldman et al., 1989; Byron and
Goldman, 1990; Schindler, 1997; Bleckner et al., 2002; Winder and
Schindler, 2004; Tranvik et al., 2009). If winter warming eliminates
ice cover altogether, hypolimnetic temperatures in the resulting
monomictic lake will parallel the annual minimum epilimnetic
temperature during the isothermal period, and will thus be more
sensitive to long-term changes in winter air temperature than in
a dimictic, seasonally ice-covered lake (Peeters et al., 2002). At SL,
we speculate that reduced or eliminated ice cover and earlier
snowmelt runoff during warm intervals results in a longer growing
season, but also a longer stratified period, potentially magnifying
the importance of processes occurring below the thermocline.
Warmer hypolimnion temperatures and stronger/longer stratification would promote increased microbial activity, more rapid OM
regeneration rates, oxygen depletion, and, perhaps most crucially,
increased microbially-mediated phosphorus (P) release from
anoxic sediments (Frevert, 1980; Carlton and Wetzel, 1988; Pace
and Prairie, 2005; Hupfer and Lewandowski, 2008; Tranvik et al.,
2009; Gudasz et al., 2010). At SL, increased supplies of regenerated DIC and DIN, combined with sediment-sourced P, would
support higher rates of summer productivity in the photic hypolimnion, the biomass from which would further fuel the microbial
ecosystem, anoxia and methanogenesis in deep waters and/or
surficial sediments. An increased supply of regenerated and
sediment-released nutrients from the hypolimnion would then be
available to support larger epilimnetic spring blooms in subsequent
years. Some similar set of mechanisms (plus an external, anthropogenic nutrient source) is consistent with trends toward lower
d13Corg, d15N and C/N, and higher TOC and BSi, in 20th century
sediments at SL. Moreover, the release of P from anoxic, high OM,
microbially-active sediments provides a plausible pathway for
nutrient enrichment during past intervals, absent anthropogenic
interference.
4.4. Late Pleistocene environmental history
4.4.1. Millennial-scale variability
4.4.1.1. High TOC intervals (SLZ-2, -4, -6, -8). The LP record was
dominated by the deposition of alternating high- and low-TOC
sediment units (Fig. 6, SLZ-2 to SLZ-9). The four high-TOC, gyttjadominated intervals of the LP (SLZ-2, 18.2e16.5 kyr BP; SLZ-4,
15.8e15.0 kyr BP; SLZ-6, 13.8e13.2 kyr BP; SLZ-8, 11.4e11.0 kyr BP)
can be interpreted as periods of increased ecosystem productivity
and reduced dilution of sedimentary TOC, reflecting drier, warmer
climate conditions and reduced sediment transport relative to SLZ-1.
The combined effects of warmer air and water temperatures, smaller
seasonal snowpacks, and the absence of glacial meltwater and till
input to the lake would have reduced winter ice cover, stimulated
lake productivity, and promoted vegetation growth and soil stabilization in the SL drainage, with the net effect of increasing OM
content and reducing detrital input to the sediments. This interpretation is consistent with the relatively high BSi and low MS
measured in the sediments of these periods. Relatively low d13Corg
(33 to 28&) (lowest values correspond to TOC maxima) may be
indicative of OM recycling and microbial production under low
oxygen conditions in the hypolimnion of a warm, stratified lake,
99
although laminated sediments occur only intermittently during
these intervals (e.g., SLZ-2, SLZ-8).
The unique combination of moderate-to-high BSi, high C/N and
low d13Corg in sediments from SLZ-2b and SLZ-4 (Figs. 4, 6 and 7)
suggests that OM derived primarily from a mixture of lake algal/
microbial and terrestrial plant sources, with reduced contributions
from aquatic plants in comparison to later periods. The long-term
trend toward lower C/N values after SLZ-4 (w15 kyr BP),
including SLZ-6 and -8, may suggest an increasing presence of
aquatic vegetation.
4.4.1.2. Low TOC intervals (SLZ-3, -5, -7, -9). The four low TOC
intervals of the LP (SLZ-3, 16.5e15.9 kyr BP; SLZ-5, 14.8e13.9 kyr BP;
SLZ-7, 13.1e11.8 kyr BP; SLZ-9, 11.0e10.7 kyr BP) each reflect lake
environments in which OM production was reduced relative to
dilution, but lithological and proxy differences among these intervals suggest that the causes of low TOC varied. The TOC minimum of
SLZ-3, consisting of massive gray clay, closely resembles SLZ-1 (i.e.,
very low TOC, BSi, and C/N; relatively high d13Corg, d15N, MS), and
likely represents a similar environment: a cold, oligotrophic lake,
with high detrital input due to sparse watershed vegetation, and,
possibly, a glacial advance back into the drainage.
Clay deposition recurred briefly during SLZ-5, but for the most
part the SLZ-5, -7 and -9 intervals consist of low OM gyttja,
including multiple laminated units. Low TOC during SLZ-5, -7, and
-9 is attributable to the combined effects of high detrital input
(high MS) and reduced lake productivity, related to high runoff
(large winter snowpacks), cold temperatures, persistent seasonal
ice cover, and low rates of OM recycling and microbial production
in the deep lake (high d13Corg). However, the sediment types
(gyttja vs. clay), less extreme TOC minima (2e6 wt%) and intermediate C/N values (mixed plant & algal sources) of SLZ-5, -7 and
-9 in comparison to SLZ-1 and -3 signal that the later low-TOC
intervals retained well-vegetated forest ecosystems and did not
involve glacial advances into the drainage. These inferences are
supported by pollen evidence for a permanent increase in the
abundance of sub-alpine conifers (and decline of alpine tundra
species) near SL after w16 kyr BP (Smith and Anderson, 1992;
Anderson, 2011) and evidence that Sierra glaciers were restricted
to high elevations (>3000 m) after 15e14 kyr BP (Clark and
Gillespie, 1997).
4.4.1.3. The Younger Dryas. SLZ-7 (13.1e11.6 kyr BP) in the SL
record coincides with the Younger Dryas (YD) climate reversal and
provides evidence for the expression of this event in the California
region. SLZ-7 consists of several distinct sediment units, suggesting
a complex structure for the YD interval at SL. A major sand/gravel
layer (658e642 cm, 13.1e12.8 kyr BP) was followed by low TOC
(2e6 wt%) laminated and massive gyttja (12.8e12.5 kyr BP), high
TOC peat (8e10 wt%, 12.5e12.2 kyr BP), and interbedded clay/
gyttja/sand (12.1e11.6 kyr BP) (Figs. 5and 6). The gyttja interval
closely resembles previous low TOC intervals (with low BSi, high C/
N, d13Corg, MS), reflecting low lake productivity and elevated
detrital input under cool, high runoff conditions (x4.4.1.2.); the
preceding sand layer may represent a series of extreme runoff
events. The peat interval, in contrast, reflects low lake levels. High
TOC and BSi, along with low MS and d13Corg, are consistent with low
detrital input and/or increased productivity and OM recycling, and
may reflect warmer, low-runoff conditions (x4.4.1.1.). Our interpretation of the early YD (13.1e12.5 kyr BP) at SL agrees with
previous evidence for depressed lake productivity and cold, wet
conditions in the Sierra Nevada region at the height of the YD
(Benson et al., 1998, 2010; Mensing, 2001; Porinchu et al., 2003;
Briggs et al., 2005; MacDonald et al., 2008; Oster et al., 2009). Proxy
patterns at SL after w12.5 kyr BP are also consistent with regional
Author's personal copy
100
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
evidence for partial warming and drying in the late YD (MacDonald
et al., 2008; Oster et al., 2009).
4.4.2. Regional climate & links to oceanographic changes
LP fluctuations in the SL OM proxy records may correspond
with climate changes inferred from Sierra Nevada studies (Fig. 9),
though dating uncertainties among the records (w100e600 yrs)
make this exercise speculative. Nonetheless, the high degree of
climatic coherence in the modern Sierra (e.g., Aguado, 1990;
Cayan and Peterson, 1993) makes it plausible to suggest a similar
level of past coherence. High TOC (warm/low-runoff) intervals in
the SL record centered at 16.9, 15.2,13.5 and 12.5 kyr BP (SLZ-2, 4,
6; SLZ-7 peat) may correspond to lake level low stands at Owens
Lake (w17.5e16.0, 15.2, 13.9, 13.2, 12.5 kyr BP) (Fig. 9d) (Benson
et al., 1997; Bacon et al., 2006), and to dry/warm conditions at
Moaning Cavern in the western Sierra foothills (15.1e14.3,
13.9e13.5, 13.0e12.5 kyr BP, Fig. 9c) (Oster et al., 2009).
Conversely, low TOC (cool/high runoff) intervals at SL centered at
w16.3, 14.4 and 12.6 kyr BP (SLZ-1, 3, 5, 7) may match Owens Lake
high stands (w15.8e15.4, 14.1, 13.0e12.5, 11.9e11.4 kyr BP;
Fig. 9d) and wet/cool intervals at Moaning Cavern (16.3e15.3,
14.2, and 12.3e11.6 kyr BP, Fig. 9c) and Pyramid Lake
(w15.5 kyr BP; 13.0e11.9 kyr BP, not shown) (Benson et al., 2010).
Previous studies noted synchronous climate changes in the North
Pacific and North Atlantic regions (e.g., GISP2 ice core record,
Fig. 9e) since Termination 1A (Bolling warming) (Benson et al.,
1997; Mensing, 2001; Barron et al., 2003; Oster et al., 2009),
and suggested a teleconnection driven by shifts in the mean
latitude of the polar jet stream (PJS) and the Pacific winter storm
track in response to the expansion and contraction of the Laurentide Ice Sheet and Arctic sea ice (Sewall and Sloan, 2004; Kim
et al., 2008). However, the SL record provides new evidence that
climate changes in the Sierra Nevada predated Greenland warming by 2000e3000 years (Fig. 9), and that factors other than ice
sheet dynamics were important determinants of regional climate
during this time.
The SL TOC record (Fig. 9a) resembles SST reconstructions from
the central and northern California margin (ODP 1017, Seki et al.,
2002; ODP 1019, Barron et al., 2003) (Fig. 1a; 9b), suggestive of
persistent links between ocean conditions in the California
Current system and Sierra Nevada climate during the LP. Within
age uncertainties (up to 630 yr in 1019 age model), each major
TOC shift at SL can be matched to a SST shift, such that high (low)
TOC corresponded to relatively warm (cold) SST off California. A
connection between coastal SST and the inland Sierra Nevada may
be direct (i.e., buffering of continental temperature by the adjacent ocean) (Herbert et al., 2001) or indirect, mediated by cooccurring changes in the ocean-atmospheric circulation of the
North Pacific. Modern studies have linked winter warming and
early snowmelt in the Sierra Nevada since 1950 to a long-term
increase in the incidence of deep low-pressure systems over the
central North Pacific (strengthened Aleutian Low), leading to
a northward displacement of the westerlies over North America,
a higher frequency of warmer air masses reaching California
Fig. 9. Late Pleistocene record of (a) sedimentary TOC at Swamp Lake, in comparison with (b) alkenone-derived SST reconstructions from ODP sites 1017 (Seki et al., 2002) and 1019
(Barron et al., 2003) off the California coast, (c) speleothem calcite d13C, a hydrologic proxy, from Moaning Cavern in the central Sierra foothills (Oster et al., 2009), (d) lake level
reconstruction for Owens Lake, adapted from Bacon et al. (2006), and (e) d18O in the Greenland GISP2 ice core (Grootes and Stuiver, 1997). Gray shading highlights local maxima in
proxy values (a, b, c, e) and low reconstructed lake level at Owens Lk. (d). SL sediment zones discussed in the text are labeled (SLZ-1 to -9), and suggested correlations in the other
records identified by a corresponding number. Age errors for the ODP 1019 and Moaning Cavern records are estimated as the mean error (calendar yrs) of all radiocarbon ages
between 10 and 20 kyr BP, from Barron et al. (2003) and Oster et al. (2009), respectively.
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
during the winter (JaneMar), and, in parallel, warmer winter SSTs
along the California margin (Dettinger and Cayan, 1995; Mantua
et al., 1997; Robertson and Ghil, 1999; Higgins et al., 2002). The
co-occurring changes at SL and California margin sites during the
LP suggests some analogous North Pacific response to outside
forcing, affecting the partitioning of sub-tropical and polar air
masses reaching the central Sierra Nevada during the winter
season, altering both temperature and precipitation regimes. The
inferred warming at SLZ-2 (w18.2 kyr BP) is consistent with
observations of pre-Bolling warming in the North Pacific (Herbert
et al., 2001; Seki et al., 2002; Kiefer and Kienast, 2005; Hill et al.,
2006; Winograd et al., 2006), and the idea that tropical Pacific
forcing contributed to patterns of deglacial warming in the
extratropics (e.g., Lea et al., 2000; Kiefer and Kienast, 2005;
Winograd et al., 2006).
4.5. Holocene environmental history at Swamp Lake
Relative to the LP, the Holocene at SL (<10.7 kyr BP) was characterized by high lake productivity and/or low detrital dilution, and
increasingly autochthonous (algal, microbial, macrophytic) OM
sources, as indicated by high sedimentary TOC and BSi, low C/N,
d13Corg and MS, and the predominance of varved gyttja throughout
the interval. Varved sediments are suggestive of anoxic conditions
at the sedimentewater interface, consistent with high lake
productivity and enhanced microbial reworking of OM and
secondary productivity (low/declining d13Corg, d15N). These conditions are consistent with a warmer, dryer mean climate relative to
the LP. Previous studies have estimated LP-Holocene warming of
3e5 C at Sierra Nevada sites (Smith and Anderson, 1992; Mensing,
101
2001; Porinchu et al., 2003) supporting our contention that sustained high TOC in Holocene sediments broadly reflects temperature regime shift. Though millennial-scale fluctuations of the sort
that characterized the LP were notably muted, the proxy records
contained several abrupt shifts (x4.5.1) and a high degree of
centennial-scale variability (x4.5.2).
4.5.1. Holocene regime shifts
Average TOC in Holocene sediments increased in three steps
(each >2 wt% C), after w10.8, 8.0 and 3.0 kyr BP (Fig. 10a), with
intervening periods characterized by stable long-term means. At
each juncture BSi also increased, while C/N and d13Corg decreased,
a pattern consistent with increasing lake productivity and
autochthonous OM, declining detrital input, and, as outlined in
x4.2.2., an increasing influence of OM regeneration and secondary
heterotrophic production in the hypolimnion on overall TOC
abundance. We use these TOC increases to divide the Holocene
record into early (EH, 10.5e8.0 kyr BP), middle (MH, 7.5e3.1 kyr BP)
and late Holocene (LH, 3.0e0 kyr BP) sub-intervals, and argue that
they were related to changes in cold-season temperature and
hydrologic regimes, influencing detrital input and OM production
and cycling within the lake, with the net effect of increasing sedimentary OM abundance.
4.5.1.1. Early Holocene (EH). Under this interpretive framework,
comparatively lower TOC and BSi, and higher average MS, C/N and
d13Corg in EH sediments are indicative of higher vascular plant OM
and detrital input relative to subsequent periods, potentially
attributable to cooler, wetter conditions. However, previous studies
at Sierra sites describe the EH as a period of warm summers, low
Fig. 10. Holocene records of sedimentary (a) TOC and (b) d13Corg at Swamp Lake, in comparison with alkenone-derived SST reconstructions from (c) ODP site 1019 (Barron et al.,
2003) and (d) ODP site 1017 (Seki et al., 2002) off Northern California. Early (EH), middle (MH) and late (LH) transitions are indicated. TOC maxima (d13Corg minima) are highlighted
in gray; low TOC, high d13Corg events are indicated by letters (aes). Suggested corresponding events in the SL and ODP 1019 records are indicated by lettering. Age error for the ODP
1019 record is estimated as the mean error (cal. yrs) of all radiocarbon ages between 0 and 10 kyr BP (Barron et al., 2003).
Author's personal copy
102
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
effective moisture and high fire frequency, supporting open forests
and extensive dryland/chaparral vegetation, associated with the
northern hemisphere summer insolation maximum (winter
minimum) at w10 kyr BP (e.g., Anderson, 1990; Smith and
Anderson, 1992; Anderson and Smith, 1994; Koehler and
Anderson, 1995; Mensing, 2001; Brunelle and Anderson, 2003;
Oster et al., 2009). A transition to warm-summer, seasonally-dry
conditions in the SL drainage after w10.8 kyr BP is also supported
by pollen evidence for the arrival of alder and near-elimination of
giant sequoia, fir and mountain hemlock (Anderson, 2011). If, as we
have argued, OM production and cycling at SL are particularly
sensitive to winter conditions (x4.3.3.2.), the observed EH proxy
values may reflect the combined effects of a highly seasonal
climate: Cold winters, with increased ice cover and a delayed
phytoplankton growing season, suppressed lake productivity and
OM cycling, while arid summers reduced vegetation density and
abetted detrital transport during spring runoff.
4.5.1.2. Middle Holocene (MH). The shift to higher baseline TOC
after w8.0 kyr BP was accompanied by increased BSi, lower C/N and
d13Corg, and a permanent lowering of MS in SL sediments, together
indicative of increased lake productivity, decreased detrital input,
and an expansion of OM recycling and heterotrophic production.
These patterns are consistent with a shift toward a drier climate
and warmer winters relative to the EH. Evidence for a regional
environmental transition w8.0 kyr BP has been found in declining
lake levels in basins draining the Sierra (Davis, 1999a; Benson et al.,
2002; Bacon et al., 2006; Negrini et al., 2006) and the expansion/
persistence of drought-tolerant vegetation on either side of the
range (e.g., Anderson, 1990; Smith and Anderson, 1992; Davis,
1999a; Brunelle and Anderson, 2003; Mensing et al., 2004;
Anderson, 2011). At SL, the period w7.4e5.5 kyr BP appears as
a broad BSi maximum (C/N minimum), indicative of increased algal
(diatom) contributions to the OM pool relative to vascular plants.
We speculate that these patterns represent an extended period of
reduced runoff during the driest interval of the MH
(w7.5e6.0 kyr BP; Lindstrom, 1990; Benson et al., 2002; Mensing
et al., 2004), when the lake was often restricted to its deep
central basin, limiting the areal extent of the shallow water habitats
utilized by aquatic/emergent vegetation (Fig. 1c). Between w5.5
and 3.2 kyr BP, BSi declined and C/N increased gradually, reflecting
reduced diatom productivity and an increased contribution of
vascular plant material to sedimentary OM in response to
increasing precipitation, runoff and lake levels, as has been inferred
from other Sierra reconstructions (Smith and Anderson, 1992;
Benson et al., 2002; Brunelle and Anderson, 2003; Mensing et al.,
2004; Osleger et al., 2009).
4.5.1.3. Late Holocene (LH). After w3.1 kyr BP, a major environmental shift occurred that over a few centuries led to large
increases in TOC and BSi and declines in C/N, d13Corg and d15N in SL
sediments. Since detrital input remained low (low MS) across the
MHeLH boundary, these patterns represent a change in the trophic
status of the lake, with increased lake primary productivity and
expanded OM recycling and secondary production. The period of
highest ecosystem productivity (high TOC, BSi) persisted until
w1.8 kyr BP, followed by a lower baseline lasting until the LIA
interval (w0.5 kyr BP). The MHeLH transition roughly corresponded to a series of regional environmental changes between
w4.0 and 3.0 kyr BP. These included glacial advances (Konrad and
Clark, 1998) and treeline retreat in the southern Sierra Nevada
(Scuderi, 1987) and White Mountains (La Marche, 1973), lake level
increases and high stands in Sierra drainages (Stine, 1990; Davis,
1999a, b; Benson et al., 2002; Mensing et al., 2004; Bacon et al.,
2006; Negrini et al., 2006), and increased effective moisture in
western Sierra meadows and forests (Anderson, 1990; Smith and
Anderson, 1992; Anderson and Smith, 1994; Brunelle and
Anderson, 2003), consistent with a regional shift toward cooler
summers and wetter winters. The evidence from SL suggests that
increasing winter temperatures may also have been an important
part of this environmental transition, inducing changes in winter
ice cover, lake mixing regime, lake productivity and OM recycling
via feedbacks similar to those outlined in x4.3.3.2. Meanwhile,
higher mean lake levels in response to a higher precipitation/
evaporation balance may have allowed for the development of the
shallow-water and marsh habitats that characterize the modern
lake (Fig. 1c; Street et al., 2008; Anderson, 2011).
4.5.2. Holocene centennial variability
Beginning w11.3 kyr BP, centennial-scale fluctuations in the SL
proxy records (events aes, Figs. 6and 10), occurring within the
longer-term regimes discussed above, became more prevalent.
Though generally smaller in magnitude than LP millennial-scale
events, these fluctuations significantly exceed measurement error
(6% of signal for >10 wt% C), encompass several data points, and
occur in multiple proxies (e.g. TOC, TN, d13Corg), implying a common
environmental cause. The apparent shift in the dominant timescale
of variability may reflect a change in the pacing of climate forcing in
the Sierra Nevada, or an increase in the sensitivity of the SL system.
Low TOC (low BSi, high C/N, high d13Corg) excursions represent
intervals of increased detrital input and/or reduced ecosystem
productivity and OM recycling. High TOC events likely reflect
relatively eutrophic conditions with reduced detrital input. The
more pronounced fluctuations observed in d13Corg compared to C/N
or d15N (Fig. 6and 10) suggest that the variability is linked to
hypolimnetic OM production and carbon cycling. Multiple low TOC
events during the EH and MH corresponded to coarse or massive
sediment intervals, or small MS peaks (“s”, 10.0 kyr BP; “r”,
9.4 kyr BP, “q”, 8.9 kyr BP; “o”, w8.0 kyr BP; “n”, 7.5 kyr BP; “k”,
5.5 kyr BP; “j”, 4.9 kyr BP; “h”, w3.7 kyr BP; “f”, 3.0 kyr BP) possibly
representing high-runoff events and/or alleviation of lake bottom
anoxia. Reduced vegetation cover or lower lake levels relative to the
present may have facilitated sediment transport to the lake depocenter during such events.
The large excursion at the EHeMH transition (7.8e7.6 kyr BP;
Fig. 6 and 10) consists of a peat layer with high TOC, C/N, d13Corg,
and d15N values e a combination unique in the SL Holocene record.
We speculate that the onset of very dry conditions at the EHeMH
transition resulted in a drastic lowering of lake level, allowing
aquatic plants (high C/N, d13Corg, Fig. 4) to colonize the lake shallows at high densities, in close proximity to the depocenter. High
abundances of floating aquatic plants can be found at present in the
shallow western end of SL, and covering several small, shallow
lakes nearby.
The last thousand years in the SL record encompass large fluctuations in OM proxies that coincide with the relatively wellcharacterized Medieval and LIA intervals. High productivity and
active OM cycling (high TOC, BSi, low d13Corg, C/N) at SL between
w1.0 and 0.5 kyr BP corresponded to generally warm and dry
Medieval climate in the Sierra Nevada (e.g., Stine, 1990, 1994;
Graumlich, 1993; Meko et al., 2001; Benson et al., 2002; Brunelle
and Anderson, 2003; Mensing et al., 2004; Yuan et al., 2004;
Hallett and Anderson, 2010). After w0.55 kyr BP, declining TOC
(BSi) and increasing C/N (d13Corg, d15N) suggest a reduction in lake
productivity and a relative increase in OM contributions from
vascular plants, in conjunction with the onset of colder temperatures, glacial advances, treeline retreat, and reduced drought
frequency in the Sierra during the LIA (e.g., Graumlich, 1993; Clark
and Gillespie, 1997; Lloyd and Graumlich, 1997; Konrad and Clark,
1998). Deposition of massive gyttja between 0.4 and 0.1 kyr BP
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
(event “a”) may represent an intensification of these trends, and an
alleviation of hypolimnetic anoxia, or a more rapid sedimentation
event (Roach, 2010). SL diatom assemblages during the LIA interval
are indicative of high lake levels and increased seasonal ice cover
(Starratt and Anderson, in press). The consistent patterns in the
proxy records over these recent, relatively well-constrained
climatic intervals bolster our interpretation of Holocene centennial variability as primarily representing the response of the lake
environment to alternating climate states.
4.5.3. Relation to California margin SST and North Pacific
circulation during the Holocene
The major Holocene baseline increases in TOC after w10.8, 8.0
and 3.0 kyr BP can be matched to corresponding shifts in reconstructed California Current SST at Site 1019, within age uncertainties
(Holocene mean 270 yr for ODP 1019; 36e188 yr for SL) (Fig. 10).
Some century-scale fluctuations in TOC at SL may match specific
events in the SST record (Fig. 10, events c, e, f, h, i, leo, qes). During
the transitional period following the YD (w12.0e10.0 kyr BP), highamplitude SST fluctuations (3 C) corresponded to the transitions
between SLZ-7, -8, -9 and the EH in the TOC record. Following the
SST/TOC peak at w10.0 kyr BP, a series of smaller, century-scale
fluctuations in each record may also correspond (Fig. 10, events
nes). Warm reconstructed SST during the EH may be attributable to
a mid-late summer weakening of the North Pacific High, slowing of
the California current, and increased penetration of warm subtropical gyre waters into the coastal zone persisting into the
fallewinter season (Barron et al., 2003; Barron and Bukry, 2007),
perhaps suggesting a cold-season link between coastal SST patterns
and TOC abundance patterns at SL.
Between 8.3 and 7.8 kyr BP (ODP 1019 age model), SST fluctuated markedly (1.3 C), before settling 1e2 C below the EH
baseline for the next w4 kyrs. This was interpreted as a general
broadening of the California current and extension of the upwelling
season, excluding gyre waters and reducing fall and winter SST
(Barron et al., 2003; Barron and Bukry, 2007). Barron and Anderson
(2011) attributed these oceanographic changes to the establishment of a North Pacific circulation regime analogous to the negative
phase of the PDO, with a weak or west-trending winter Aleutian
Low and a strong summer North Pacific High. Within the MH
interval (w8.0e3.0 ka), several century-scale TOC and SST fluctuations may correspond (Fig. 10) though the correlations become
more tentative after w6.0 kyr BP.
Between 3.4 and 3.1 kyr BP (1019 age model), alkenone SST
increased by w1.5 C, attributable to increased fall and winter SST
related to a seasonal weakening of the California current (Barron
et al., 2003; Barron and Bukry, 2007). This regime shift, and the
subsequent SST optimum (to w2 kyr BP), are likely correlative with
the large increase and sustained TOC maximum in the SL record
occurring after w3.0 kyr BP (3.0e1.8 kyr BP), which we attribute
largely to higher winterespring temperatures. LH oceanographic
changes are consistent with a North Pacific regime-shift towards an
El Nino/þPDO-like state, characterized by a strong and/or
eastward-shifted winter Aleutian Low, a weakened North Pacific
High, and anomalously warm winterespring temperatures and
somewhat elevated precipitation in the Sierra Nevada (Zhang et al.,
1997; Gershunov and Barnett, 1998; Cayan et al., 2001; Mantua and
Hare, 2002; Barron et al., 2003; Castello and Shelton, 2004; Barron
and Anderson, 2011). Warmer winterespring conditions at SL after
3.0 kyr BP would likely have stimulated positive feedback loops in
the lake ecosystem, ultimately resulting in the observed LH TOC
increase (x4.3.3.2.; x4.5.1.3.).
In summary, most large changes in the abundance and
composition of OM in SL sediments over the past 18,000 years were
associated with SST shifts in the California current system off the
103
northern and central California coast. This relationship was most
noticeable during the millennial TOC fluctuations of the LP and
major transitions of the Holocene, but may also be discernable
during century-scale events of the Holocene. The relationship
probably reflects large-scale changes in North Pacific oceanatmospheric circulation, which altered temperature and precipitation regimes (especially during winter) in the Sierra Nevada in
parallel with surface ocean conditions in the California current. The
major exception to the generally positive SSTeTOC relationship
occurred at the EHeMH transition (w8.0e7.5 kyr BP), when baseline TOC at SL increased while SST declined. Here, the local
hydrologic effects of this climate transition (reduced runoff and
detrital input, lower lake levels and changed mixing regimes) may
have outweighed other influences. The abrupt increases in TOC at
SL and SST off California at the MHeLH transition are broadly
consistent with a late Holocene intensification of ENSO cycling in
tropical Pacific records (e.g., Moy et al., 2002; Conroy et al., 2008)
and a shift toward a þPDO-like mean state in the North Pacific
circulation (Barron and Anderson, 2011).
5. Conclusions
The Swamp Lake OM record provides new insight into the
patterns and pacing of environmental change in the Sierra Nevada
since the LGM. Variability in sedimentary TOC abundance was
closely associated with changes in OM composition (C/N, d13Corg,
d15N), BSi concentration, total magnetic susceptibility, and sedimentology. These observations indicate that TOC over time has
been governed by changes in lake total productivity (including both
algal and heterotrophic sources), vascular plant OM inputs, and
dilution by detrital materials eroded from the surrounding
drainage. These factors in turn responded to changes in lake
temperature, winter
ice
cover, mixing
regimes,
and
springesummer runoff in response to shifting climate conditions.
Very low sedimentary d13Corg values, particular during high TOC,
varved sediment intervals, combined with modern water column
measurements (temperature, chl a, DO, d13CPOM) suggest that the
intensity of OM regeneration and secondary microbial productivity
under low-oxygen conditions in the hypolimnion and surficial
sediments is an important control on sedimentary TOC.
The late Pleistocene (LP) record (19.7e10.7 kyr BP) was dominated by millennial-scale fluctuations between high and low TOC
sedimentation, reflecting major changes in regional climate and
local responses in the SL basin. Low TOC intervals (19.7e18.2,
16.5e15.8 kyr BP; 14.9e13.9 kyr BP; 13.1e11.7 kyr BP;
11.0e10.7 kyr BP), at times associated with increased deposition
of clastic sediments, reflect oligotrophic conditions and heavy
detrital input to the lake basin under cold, wet climates. Clay and
silt deposition during earlier low TOC periods (19.7e18.2,
16.5e15.8 kyr BP) was related to erosion from poorly-vegetated,
post-glacial landscapes; establishment of a sub-alpine forest
ecosystem (by w15.0 kyr BP) resulted in a more stable watershed
and less distinctive sedimentation during later low TOC intervals.
High TOC gyttja intervals (17.4e16.5, 15.8e15.0, 13.9e13.2,
11.4e11.0 kyr BP), in contrast, are associated with reduced detrital
dilution, increased lake productivity and plant OM inputs, and
active deep-lake OM cycling under relatively warm/dry climate
conditions.
After w12 kyr BP, SL sediments were predominantly varved,
reflecting persistent seasonal anoxia in the deep lake and, after
w10.5 kyr BP, sustained high productivity conditions. The Holocene
record was defined by three abrupt increases in TOC after w10.8,
8.0 and 3.0 kyr BP, which we attribute to increases in lake
productivity, deep lake OM cycling and secondary production and/
or decreased detrital dilution and allocthonous OM input, related to
Author's personal copy
104
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
warmer winterespring temperatures and altered hydrologic
regimes. Numerous centennial-scale fluctuations between warm/
dry (high TOC) and cool/wet (low TOC) conditions also occurred
within the early, middle, and late Holocene regimes.
Major changes in TOC and other proxies at SL on both millennial
and centennial timescales corresponded to changing SST regimes in
the California current system. The coherence between these
records, persisting across all major climate transitions since the
LGM, suggests a common forcing mechanism arising in the oceanatmospheric circulation of the North Pacific.
Acknowledgments
Three anonymous referees contributed reviews that improved
the quality of the paper. Scott Starratt, Douglas Hallett and John
Barron also provided helpful comments at various points in this
project. We thank Jan van Wagtendonk, Dan Cayan, John Barron
and James Stringfellow for logistical assistance; Daniel Boone, Jordon Bright, Alison Colwell, Elizabeth Derse, Douglas Hallett, Justin
Holl, Erin Hult, Karen Knee, Tim Lambert, Jackie Liu, Ian McKenna,
Nadine Quintana-Krupinski, Lydia Roach, and Benjamin Saenz for
field assistance; and Cara Meeker and Susan Smith for laboratory
assistance. Initial core characterization was carried out at LacCore
(National Lacustrine Core Facility), Department of Geology and
Geophysics, University of Minnesota. Biogenic silica measurements
were made by Jordon Bright in the Sedimentary Records of Environmental Change laboratory at NAU. This research was supported
by grants from the National Science Foundation (EAR-0902218)
(A.P.), Calfed Bay-Delta Science Program (J.S.), and The Yosemite
Fund (R.S.A.), and by funding from the U.S. Geological Survey. J.S.
was supported by graduate student fellowships from Stanford
University and the Calfed Bay-Delta Science Program. Laboratory of
Paleoecology Contribution # 132.
References
Aguado, E., 1990. Elevational and latitudinal patterns of snow accumulation
departures from normal in the Sierra Nevada. Theor. Appl. Climatol. 42,
177e185.
Anderson, R.S., 1990. Holocene forest development and paleoclimates within the
central Sierra Nevada, California. J. Ecol. 78, 470e489.
Anderson, R.S., 2011. Fire and Vegetation History from Swamp Lake, Yosemite
National Park, California: A Final Report on Research. Technical Report, The
Yosemite Conservancy, San Francisco, California, USA.
Anderson, R.S., Smith, S.J., 1994. Paleoclimatic interpretations of meadow sediment
and pollen stratigraphies from California. Geology 22, 723e726.
Bacon, S.N., Burke, R.M., Pezzopane, S.K., Jayko, A.S., 2006. Last glacial maximum
and Holocene lake levels of Owens Lake, eastern California, USA. Quat. Sci. Rev.
25, 1264e1282.
Barnola, J.-M., Raynaud, D., Lorius, C., Barkov, N.I., 2003. Historical CO2 record from
the Vostok Ice Core. In: Trends: A Compendium of Data on Global Change.
Carbon Dioxide Information Analysis Center, Oak Ridge National Laboratory,
U.S. Department of Energy, Oak Ridge, Tenn., U.S.A.
Barron, J.A., Bukry, D., 2007. Development of the California Current during the past
12,000-yr based on diatoms and silicoflagellates. Palaeogeogr., Palaeoclimatol.,
Palaeoecol. 248, 313e338.
Barron, J.A., Anderson, L., 2011. Enhanced Late Holocene ENSO/PDO expression
along the margins of the eastern North Pacific. Quat. Int. 235, 3e12.
Barron, J.A., Heusser, L., Herbert, T., Lyle, M., 2003. High-resolution climatic evolution of coastal northern California during the past 16,000 years. Paleoceanography 18, 1020. doi:10.1029/2002PA000768.
Bastviken, D., Ejlertsson, J., Sundh, I., Tranvik, L., 2003. Methane as a source of
carbon and energy for lake pelagic food webs. Ecology 84, 969e981.
Benson, L.V., Burdett, J.W., Kashgarian, M., Lund, S.P., Phillips, F.M., Rye, R.O., 1996.
Climatic and hydrologic oscillations in the Owens Lake Basin and adjacent
Sierra Nevada, California. Science 274, 746e749.
Benson, L., Burdett, J., Lund, S., Kashgarian, M., Mensing, S., 1997. Nearly synchronous climate change in the Northern Hemisphere during the last glacial
termination. Nature 388, 263e265.
Benson, L.V., May, H.M., Antweiler, R.C., Brinton, T.I., Kashgarian, M., Smoot, J.P.,
Lund, S.P., 1998. Continuous lake-sediment records of glaciation in the Sierra
Nevada between 52,600 and 12,500 14C yr B.P. Quat. Res. 50, 113e127.
Benson, L., Kashgarian, M., Rye, R., Lund, S., Paillet, F., Smoot, J., Kester, C.,
Mensing, S., Meko, D., Lindstrom, S., 2002. Holocene multidecadal and
multicentennial droughts affecting Northern California and Nevada. Quat. Sci.
Rev. 21, 659e682.
Benson, L., Linsley, B., Smoot, J., Mensing, S., Lund, S., Stine, S., Sarna-Wojcicki, A.,
2003. Influence of the Pacific Decadal Oscillation on the climate of the Sierra
Nevada, California and Nevada. Quat. Res. 59, 151e159.
Benson, L.V., Lund, S., Smoot, J., 2010. A Comparison of Climate Change in the
Lahontan and Bonneville Basins with the GISP2 d18O Record for the Period
45e11 cal ka. Abstract GC23H-01 presented at 2010 Fall Meeting, AGU, San
Francisco, Calif., 13e17 Dec.
Biondi, F., Gershunov, A., Cayan, D.R., 2001. North Pacific decadal climate variability
since 1661. J. Climate 14, 5e10.
Bleckner, T., Omstedt, A., Rummukainen, M., 2002. A Swedish case study of
contemporary and possible future consequences of climate change on lake
function. Aquat. Sci. 64, 171e184.
Brenner, M., Whitmore, T.J., Curtis, J.H., Hodell, D.A., Schelske, C.L., 1999. Stable
isotope (d13C and d15N) signatures of sedimented organic matter as indicators of
historic lake trophic state. J. Paleolimnol. 22, 205e221.
Briggs, R.W., Wesnousky, S.G., Adams, K.D., 2005. Late Pleistocene and late Holocene
lake highstands in the Pyramid Lake subbasin of Lake Lahontan, Nevada, USA.
Quat. Res. 64, 257e263.
Brunelle, A., Anderson, R.S., 2003. Sedimentary charcoal as an indicator of lateHolocene drought in the Sierra Nevada, California, and its relevance to the
future. Holocene 13, 21e28.
Byron, E.R., Goldman, C.R., 1990. The potential effects of global warming on the
primary productivity of a subalpine lake. Water Res. Bull. 26, 983e989.
Carlton, R.G., Wetzel, R.G., 1988. Phosphorus flux from lake sediments: effect of
epipelic algal oxygen production. Limnol. Oceanogr. 33, 562e570.
Castello, A.F., Shelton, M.F., 2004. Winter precipitation on the US Pacific coast and El
Nino-Southern Oscillation events. Int. J. Climatol. 24, 481e497.
Cayan, D.R., Peterson, D.H., 1993. Spring climate and salinity in the San Francisco
Bay Estuary. Water Resour. Res. 29, 293e303.
Cayan, D.R., Kammerdiener, S.A., Dettinger, M.D., Caprio, J.M., Peterson, D.H., 2001.
Changes in the onset of spring in the western United States. B. Am. Meteorol.
Soc. 82, 399e415.
Clark, D.H., Gillespie, A.R., 1997. Timing and significance of late-glacial and Holocene
cirque glaciation in the Sierra Nevada, California. Quat. Int. 38/39, 21e38.
Clow, D.W., Sickman, J.O., Striegl, R.G., Krabbenhoft, D.P., Elliott, J.G., Dornblaser, M.,
Roth, D.A., Campbell, D.H., 2003. Changes in the chemistry of lakes and
precipitation in high-elevation national parks in the western United States,
1985e1999. Water Resour. Res. 39, 1171. doi:10.1029/2002WR001533.
Colman, S.M., Peck, J.A., Karabanov, K.A., Carter, S.J., Bardbury, J.P., King, J.W.,
Williams, D.F., 1995. Continental climate response to orbital forcing from
biogenic silica records in Lake Baikal. Nature 378, 769e771.
Conroy, J.L., Overpeck, J.T., Cole, J.E., Shanahan, T.M., Steinitz-Kannan, M., 2008.
Holocene changes in eastern tropical Pacific climate inferred from a Galápagos
lake sediment record. Quat. Sci. Rev. 27, 1166e1180.
Conroy, J.L., Overpeck, J.T., Cole, J.E., Steinitz-Kannan, M., 2009. Variable oceanic
influences on western North American drought over the last 1200 years. Geophys. Res. Lett. 36, L17703. doi:10.1029/2009GL039558.
Cook, E.R., Woodhouse, C.A., Eakin, C.M., Meko, D.M., Stahle, D.W., 2004. Long-term
aridity changes in the western United States. Science 306, 1015e1018.
Davis, O.K., 1999a. Pollen analysis of a late-glacial and Holocene sediment core from
Mono Lake, Mono County, California. Quat. Res. 52, 243e249.
Davis, O.K., 1999b. Pollen analysis of Tulare Lake, California: Great Basin-like
vegetation in Central California during the full-glacial and early Holocene.
Rev. Palaeobotany and Palynology 107, 249e257.
Davis, O.K., Moratto, M.J., 1988. Evidence for a warm dry early Holocene in the
western Sierra Nevada of California: pollen and plant macrofossil analysis of
Dinkey and Exchequer Meadows. Madroño 35, 132e149.
Davis, O.K., Anderson, R.S., Fall, P.L., O’Rouke, M.K., Thompson, R.S., 1985. Palynological evidence for early Holocene aridity in the southern Sierra Nevada, California. Quat. Res. 24, 322e332.
Dettinger, M.D., Cayan, D.R., 1995. Large-scale atmospheric forcing of recent trends
toward early snowmelt runoff in California. J. Climate 8, 606e623.
Dettinger, M.D., Battisti, D.S., Garreaud, R.D., McCable, G.J., Bitz, C.M., 2001. Interhemispheric effects of interannual and decadal ENSO-like climate variations on the
Americas. In: Markgraf, V. (Ed.), Interhemispheric Climate Linkages: Present and
Past Climates in the Americas and Their Societal Effects. Academic Press, pp. 1e16.
Diaz, H.F., Hoerling, M.P., Eischeid, J.K., 2001. ENSO variability, teleconnections and
climate change. Int. J. Climatol. 21, 1845e1862.
Frevert, T., 1980. Dissolved oxygen dependent phosphorus release from profundal
sediments of Lake Constance (Obersee). Hydrobiol. 74, 17e28.
Galman, V., Rydberg, J., Sjostedt de-Luna, S., Bindler, R., Renberg, I., 2008. Carbon
and nitrogen loss rates during aging of lake sediment: changes over 27 years
studied in varved lake sediment. Limnol. Oceanogr. 53, 1076e1082.
Gershunov, A., Barnett, T.P., 1998. Interdecadal modulation of ENSO teleconnections.
B. Am. Meteorol. Soc. 79, 2715e2725.
Goldman, C.R., Jassby, A., Powell, T., 1989. Interannual fluctuations in primary production: meteorological forcing of two subalpine lakes. Limnol. Oceanogr. 34, 310e323.
Gong, C., Hollander, D.J., 1997. Differential contribution of bacteria to sedimentary
organic matter in oxic and anoxic environments, Santa Monica Basin, California.
Org. Geochem. 26, 545e563.
Graham, N.E., Hughes, M.K., Ammann, C.M., Cobb, K.M., Hoerling, M.P., Kennett, D.J.,
Kennett, J.P., Rein, B., Stott, L., Wigand, P.E., Xu, T., 2007. Tropical Pacific e midlatitude teleconnections in medieval times. Climatic Change 83, 241e285.
Author's personal copy
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
Graumlich, L.J., 1993. A 1000-year record of temperature and precipitation in the
Sierra Nevada. Quat. Res. 39, 249e255.
Grootes, P.M., Stuiver, M., 1997. Oxygen 18/16 variability in Greenland snow and ice
with 103 to 105-year time resolution. J. Geophys. Res. 102, 26455e26470.
Gu, B., Alexander, V., Schell, D.M., 1999. Seasonal and interannual variability of
plankton carbon isotope ratios in a subarctic lake. Freshwater Biol. 42, 417e426.
Gudasz, C., Bastviken, D., Steger, K., Premke, K., Sobek, S., Tranvik, L.J., 2010.
Temperature-controlled organic carbon mineralization in lake sediments.
Nature 466, 478e482.
Hallett, D.J., Hills, L.V., Clague, J.J., 1997. New accelerator mass spectrometry radiocarbon ages for the Mazama tephra layer from Kootenay National Park, British
Columbia, Canada. Can. J. Earth Sci. 34, 1202e1209.
Hallett, D.J., Anderson, R.S., 2010. Paleofire reconstruction for high-elevation forests
in the Sierra Nevada, California, with implications for wildfire synchrony and
climate variability in the late Holocene. Quat. Res. 73, 180e190.
Harris, D., Howarth, W.R., van Kessel, C., 2001. Acid fumigation of soils to remove
carbonates prior to total organic carbon or carbon-13 isotopic analysis. Soil Sci.
Soc. Am. J. 65, 1853e1856.
Herbert, T.D., Schuffert, J.D., Andreasen, D., Heusser, L., Lyle, M., Mix, A., Ravelo, A.C.,
Stott, L.D., Herguera, J.C., 2001. Collapse of the California Current during glacial
maxima linked to climate change on land. Science 293, 71e76.
Herczeg, A.L., Smith, A.K., Dighton, J.C., 2001. A 120 year record of changes in
nitrogen and carbon cycling in Lake Alexandrina, South Australia: C: N, d15N and
d13C in sediments. Appl. Geochem. 16, 73e84.
Herweijer, C., Seager, R., Cook, E.R., Emile-Geay, J., 2007. North American droughts of
the last millennium from a gridded network of tree-ring data. J. Climate 20,
1353e1376.
Higgins, R.W., Leetmaa, A., Kousky, V.E., 2002. Relationships between climate
variability and winter temperature extremes in the United States. J. Climate 15,
1555e1572.
Hill, T.M., Kennett, J.P., Pak, D.K., Behl, R.J., Robert, C., Beaufort, L., 2006. Pre-Bølling
warming in Santa Barbara Basin, California: surface and intermediate water
records of early deglacial warmth. Quat. Sci. Rev. 25, 2835e2845.
Hodell, D.A., Schelske, C.L., 1998. Production, sedimentation, and isotopic composition of organic matter in Lake Ontario. Limnol. Oceanogr. 43, 200e214.
Hollander, D.J., McKenzie, J.A., 1991. CO2 control on carbon isotope fractionation
during photosynthesis: a paleo-p CO2 barometer. Geology 19, 929e932.
Hollander, D.J., Smith, M.A., 2001. Microbially mediated carbon cycling as a control
on the 13C of sedimentary carbon in eutrophic Lake Mendota (USA): new
models for interpreting isotopic excursions in the sedimentary record. Geochim. Cosmochim. Ac. 62, 4321e4337.
Hu, F.S., Kaufman, D., Yoneji, S., Nelson, D., Shemesh, A., Huang, Y., Tian, J., Bond, G.,
Clegg, B., Brown, T., 2003. Cyclic variation and solar forcing of Holocene climate
in the Alaskan subarctic. Science 301, 1890e1893.
Hughes, M.K., Graumlich, L.G., 1996. Climatic Variations and Forcing Mechanisms of
the Last 2000 Years. In: Multi-millenial Dendroclimatic Studies from the
Western United States, vol. 141. NATO ASI Series, 109e124 pp.
Hunter, T., Tootle, G., Piechota, T., 2006. Oceanic-atmospheric variability and
western U.S. snowfall. Geophys. Res. Lett. 33, L13706. doi:10.1029/
2006GL026600.
Hupfer, M., Lewandowski, J., 2008. Oxygen controls the phosphorus release from
lake sediments- a long lasting paradigm in limnology. Int. Rev. Hydrobiol. 93,
415e432.
Jassby, A.D., Reuter, J.E., Axler, R.P., Goldman, C.R., Hackley, S.H., 1994. Atmospheric
deposition of nitrogen and phosphorus in the annual nutrient load of Lake
Tahoe (California-Nevada). Water Resour. Res. 30, 2207e2216.
Kendall, C., 1998. Tracing nitrogen sources and cycling in catchments. In: Kendall, C.,
McDonnell, J.J. (Eds.), Isotope Tracers in Catchment Hydrology. Elsevier Science
B.V., Amsterdam, pp. 519e576.
Kiefer, T., Kienast, M., 2005. Patterns of deglacial warming in the Pacific Ocean:
a review with emphasis on the time interval of Heinrich event 1. Quat. Sci. Rev.
24, 1063e1081.
Kim, S.-J., Crowley, T.J., Erickson, D.J., Govindasamy, B., Duffy, P.B., Lee, B.Y., 2008.
High- resolution climate simulation of the last glacial maximum. Clim. Dyn. 31,
1e6.
Konrad, S.K., Clark, D.H., 1998. Evidence for an early Neoglacial glacier advance from
rock glaciers and lake sediments in the Sierra Nevada, California, U.S.A. Arctic
Alpine Res. 30, 272e284.
Koehler, P.A., Anderson, R.S., 1994. The paleoecology and stratigraphy of Nichols
Meadow, Sierra National Forest, California, USA. Palaeogeogr. Palaeoclimatol.
Palaeoecol. 112, 1e17.
Koehler, P.A., Anderson, R.S., 1995. Thirty thousand years of vegetation changes in
the Alabama Hills, Owens Valley, California. Quat. Res. 24, 322e332.
La Marche Jr., V.C., 1973. Holocene climatic variations inferred from treeline fluctuations in the White Mountains, California. Quat. Res. 3, 632e660.
Laws, E.A., Popp, B.N., Bidigare, R.R., Kennicutt, M.C., Macko, S.A., 1995. Dependence
of phytoplankton carbon isotopic composition on growth rate and [CO2aq]:
theoretical considerations and experimental results. Geochim. Cosmochim. Ac.
59, 1131e1138.
LaZerte, B.D., Szalados, J.E., 1982. Stable carbon isotope ratio of submerged freshwater macrophytes. Limnol. Oceanogr. 27, 413e418.
Lea, D.W., Pak, D.K., Spero, H.J., 2000. Climate impact of late Quaternary equatorial
Pacific sea surface temperature variations. Science 289, 1719e1724.
Lehmann, M.F., Bernasconi, S.M., Barbieri, A., McKenzie, J.A., 2002. Preservation of
organic matter and alteration of its carbon and nitrogen isotope composition
105
during simulated and in situ early diagenesis. Geochim. Cosmochim. Ac. 66,
3573e3584.
Li, L., Yu, Z., Moeller, R.E., Bebout, G.E., 2008. Complex trajectories of aquatic and
terrestrial ecosystem shifts caused by multiple human-induced environmental
stresses. Geochim. Cosmochim. Ac. 72, 4338e4351.
Lindstrom, S., 1990. Submerged tree stumps as indicators of Mid-Holocene aridity in
the Lake Tahoe basin. J. California Great Basin Anthropol. 12, 146e157.
Lloyd, A.H., Graumlich, L.J., 1997. Holocene dynamics of treeline forests in the Sierra
Nevada. Ecology 78, 1199e1210.
MacDonald, G.M., Moser, K.A., Bloom, A.M., Porinchu, D.F., Potito, A.P., Wolfe, B.B.,
Edwards, T.W.D., Petel, A., Orme, A.R., Orme, A.J., 2008. Evidence of temperature
depression and hydrological variations in the eastern Sierra Nevada during the
Younger Dryas stade. Quat. Res. 70, 131e140.
MacDonald, G.M., Case, R.A., 2005. Variations in the Pacific Decadal Oscillation over
the past millennium. Geophys. Res. Lett. 32, L08703. doi:10.1029/
2005GL022478.
Mantua, N.J., Hare, S.R., Zhang, Y., Wallace, J.M., Francis, R.C., 1997. A Pacific interdecadal climate oscillation with impacts on salmon production. B. Am. Meteorol. Soc. 78, 1069e1079.
Mantua, N.J., Hare, S.R., 2002. The Pacific Decadal Oscillation. J. Oceanogr. 58,
35e44.
McCabe, G.J., Dettinger, M.D., 2002. Primary modes and predictability of year-toyear snowpack variations in the western United States from teleconnections
with Pacific Ocean climate. J. Hydrometeorol. 3, 13e25.
McGowan, J.A., Cayan, D.R., Dorman, L.M., 1998. Climate-ocean variability and
ecosystem response in the Northeast Pacific. Science 281, 210e217.
Meko, D., Therrel, M., Baisan, C., Hughes, M., 2001. Sacramento River flow reconstructed to A.D. 869 from tree rings. J. Am. Water Resources Assoc. 37,
1029e1039.
Mensing, S.A., 2001. Late Glacial and early Holocene vegetation and climate change
near Owens Lake, Eastern California. Quat. Res. 55, 57e67.
Mensing, S.A., Benson, L.V., Kashgarian, M., Lund, S., 2004. A Holocene pollen record
of persistent droughts from Pyramid Lake, Nevada, USA. Quat. Res. 62, 29e38.
Meyers, P.A., 1994. Preservation of source identification of sedimentary organic
matter during and after deposition. Chem. Geol. 144, 289e302.
Meyers, P.A., 2003. Applications of organic geochemistry to paleolimnological
reconstructions: a summary of examples from the Laurentian Great Lakes. Org.
Geochem. 34, 261e289.
Meyers, P.A., Ishiwatari, R., 1993. Lacustrine organic geochemistry e an overview of
indicators of organic matter sources and diagenesis in lake sediments. Org.
Geochem. 20, 867e900.
Meyers, P.A., Takemura, K., 1997. Quaternary changes in delivery and accumulation
of organic matter in sediments off Lake Biwa. Japan. J. Paleolimnol. 18, 211e218.
Meyers, P.A., Lallier-Verges, E., 1999. Lacustrine sedimentary organic matter records
of Late Quaternary paleoclimates. J. Paleolimnol. 21, 345e372.
Millar, C.I., Westfall, R.D., Delany, D.L., King, J.C., Graumlich, L.J., 2004. Response of
subalpine conifers in the Sierra Nevada, California, U.S.A., to 20th-century warming
and decadal climate variability. Arctic Antarctic Alpine Res. 36, 181e200.
Miller, N.L., Bashford, K.E., Strem, E., 2003. Potential impacts of climate change on
California hydrology. J. Am. Water Resources Assoc. 39, 771e784.
Mock, C.J., Bartlein, P.J., 1995. Spatial variability of late-Quaternary paleoclimates in
the western United States. Quat. Res. 44, 425e433.
Mortlock, R.A., Froelich, P.N., 1989. A simple method for the rapid determination of
biogenic opal in pelagic marine sediments. Deep-Sea Res. 36, 1415e1426.
Moy, C.M., Seltzer, G.O., Rodbell, D.T., Anderson, D.M., 2002. Variability of El Niño/
Southern Oscillation activity at millennial timescales during the Holocene
epoch. Nature 420, 162e165.
Negrini, R.M., Wigand, P.E., Draucker, S., Gobalet, K., Gardner, J.K., Sutton, M.Q.,
Yohe II, R.M., 2006. The Rambla highstand shoreline and the Holocene lakelevel history of Tulare Lake, California, USA. Quat. Sci. Rev. 25, 1599e1618.
O’Leary, M.H., 1988. Carbon isotopes in photosynthesis. Bioscience 38, 328e336.
Osleger, D.A., Heyvaert, A.C., Stoner, J.S., Verosub, K.L., 2009. Lacustrine turbidites as
indicators of Holocene storminess and climate: Lake Tahoe, California and
Nevada. J. Paleolimnol. 42, 103e122.
Oster, J.L., Montanez, I.P., Sharp, W.D., Cooper, K.M., 2009. Late Pleistocene California
droughts during deglaciation and Arctic warming. Earth Planet. Sci. Lett. 288,
434e443.
Pace, M.L., Prairie, Y.T., 2005. Respiration in lakes. In: del Giorgio, P.A., Williams, P.J.
leB. (Eds.), Respiration in Aquatic Systems. Oxford University Press, Oxford,
pp. 103e121.
Peeters, F., Livingstone, D.F., Goudsmit, G.-H., Kipfer, R., Forster, R., 2002. Modeling
50 years of historical temperature profiles in a large central European lake.
Limnol. Oceanogr. 47, 186e197.
Porinchu, D.F., MacDonald, G.M., Bloom, A.M., Moser, K.A., 2003. Late Pleistocene
and early Holocene climate and limnological changes in the Sierra Nevada,
California, USA inferred from midges (Insecta: Diptera: Chironomidae). Palaeogeog. Palaeoclimatol. Palaeoecol. 198, 403e422.
Prahl, F.G., Ertel, J.R., Goni, M.A., Sparrow, M.A., Eversmeyer, B., 1994. Terrestrial
organic carbon contributions to sediments on the Washington margin. Geochim. Cosmochim. Ac. 58, 3055e3058.
Rau, G., 1978. Carbon-13 depletion in a subalpine lake: carbon flow implications.
Science 201, 901e902.
Redmond, K.T., Koch, R.W., 1991. Surface climate and streamflow variability in the
western United States and their relationship to large-scale circulation indices.
Water Resour. Res. 27, 2381e2399.
Author's personal copy
106
J.H. Street et al. / Quaternary Science Reviews 40 (2012) 89e106
Reimer, P.J., Ballie, M.G.L., Bard, E., Bayliss, A., Beck, J.W., Blackwell, P.G., Buck, C.E.,
Burr, G.S., Cutler, K.B., Damon, P.E., Edwards, R.L., Fairbanks, R.G., Friedrich, M.,
Guilderson, T.P., Herring, C., Hughen, K.A., Kromer, B., McCormac, F.G.,
Manning, S.W., Ramsey, C.B., Reimer, R.W., Remmele, S., Southon, J.R., Stuiver, S.,
Talamo, S., Taylor, F.W., Van der Plicht, J., Weyhenmeyer, C.E., 2004. IntCal04
terrestrial radiocarbon age calibration, 0-26 cal kyr BP. Radiocarbon 46,
1029e1058.
Roach, L.D., 2010. Climate change in the Pacific North America region over the past
millennium: development and application of novel geochemical tracers. Ph.D.
Thesis, Scripps Institution of Oceanography, University of California, San Diego.
Robertson, A.W., Ghil, M., 1999. Large-scale weather regimes and local climate over
the western United States. J. Climate 12, 1796e1813.
Schindler, D.W., 1997. Widespread effects of climate warming on freshwater
ecosystems in North America. Hydrol. Process. 11, 1043e1067.
Scuderi, L.A., 1987. Late-Holocene upper timberline variation in the southern Sierra
Nevada. Nature 325, 242e244.
Seager, R., Vecchi, G.A., 2010. Greenhouse warming and the 21st century hydroclimate of southwestern North America. Proc. Nat. Ac. Sci. 107, 21277e21282.
Seki, O., Ishiwatari, R., Matsumoto, K., 2002. Millennial climate oscillations in NE
Pacific surface waters over the last 82 kyr: new evidence from alkenones.
Geophys. Res. Lett. 29, 2144. doi:10.1029/2002GL01520.
Sewall, J.O., Sloan, L.C., 2004. Disappearing Arctic sea ice reduces available water in
the American west. Geophys. Res. Lett. 31, L06209. doi:10.1029/2003GL019133.
Sickman, J.O., Lydecker, A., Chang, C.C.Y., Kendall, C., Melack, J.M., Lucero, D.M.,
Schimel, J., 2003. Mechanisms underlying export of N from high-elevation
catchments during seasonal transitions. Biogeochem. 64, 1e24.
Smith, S.J., Anderson, R.S., 1992. Late Wisconsin paleoecologic record from Swamp
Lake, Yosemite National Park, California. Quat. Res. 38, 91e102.
Starratt, S.W., Anderson, R.S., Preliminary report on the Late Pleistocene and
Holocene diatoms of Swamp Lake, Yosemite National Park, California, USA.
Proceedings of the 21st International Diatom Symposium, in press.
Stine, S., 1990. Late Holocene fluctuations of Mono Lake, Eastern California. Paleogeogr. Paleoclimatol. Paleoecol. 78, 333e381.
Stine, S., 1994. Extreme and persistent drought in California and Patagonia during
mediaeval time. Nature 369, 546e549.
Street, J.H., Anderson, R.S., Sessions, A., Paytan, A., 2008. Mid-Holocene climate
variability in the Sierra Nevada as recorded in sediments from Swamp Lake,
Yosemite NP. Eos Trans. AGU 90 (53), Fall Meet. Suppl., Abstract PP41B-1440.
Strub, P.T., Powell, T., Goldman, C.R., 1985. Climatic forcing: effects of El Nino on
a small, temperate lake. Science 227, 55e57.
Stuvier, M., Reimer, P.J., Reimer, R., 2005. CALIB Radiocarbon Calibration 5.0.2.
http://radiocarbon.pa.qub.ac.uk.
Talbot, M.R., Laerdal, T., 2000. The late Pleistocene-Holocene paleolimnology of Lake
Victoria, East Africa, based upon elemental and isotopic analyses of sedimentary organic matter. J. Paleolimnol. 23, 141e164.
Tenzer, G.E., Meyers, P.A., Knoop, P.A., 1997. Sources and distribution of organic and
carbonate carbon in surface sediments of Pyramid Lake, Nevada. J. Sediment.
Res. 67, 887e893.
Teranes, J.L., Bernasconi, S.M., 2005. Factors controlling 13C values of sedimentary
carbon in hypertrophic Baldeggersee, Switzerland, and implications for interpreting isotope excursions in lake sedimentary records. Limnol. Oceanogr. 50,
914e922.
Tootle, G.A., Piechota, T.C., Singh, A., 2005. Coupled oceanic-atmospheric variability
and U.S. streamflow. Water Resour. Res. 41, W12408. doi:10.1029/
2005WR004381.
Tranvik, L.J., Downing, J.A., Cotner, J.B., Loiselle, S.A., Striegl, R.G., Ballatore, T.J.,
Dillon, P., Finlay, K., Fortino, K., Knoll, L.B., Kortelainen, P.L., Kutser, T., Larsen, S.,
Laurion, I., Leech, D.M., McCallister, S.L., McKnight, D.M., Melack, J.M.,
Overholt, E., Porter, J.A., Prairie, Y., Renwick, W.H., Roland, F., Sherman, B.S.,
Schindler, D.W., Sobek, S., Tremblay, A., Vanni, M.J., Verschoor, A.M., von
Wachenfeldt, E., Weyhenmeyer, G.A., 2009. Lakes and reservoirs as regulators of
carbon cycling and climate. Limnol. Oceanogr. 54, 2298e2314.
Whiticar, M.J., 1999. Carbon and hydrogen isotope systematics of bacterial formation and oxidation of methane. Chem. Geol. 161, 291e314.
Wilkinson, R., Clarke, K., Goodchild, M., Reichman, J., Dozier, J., 2002. The Potential
Consequences of Climate Variability and Change: The California Regional
Assessment, A Report of the California Regional Assessment Group for the U.S.
Global Change Research Program, 432 pp.
Winder, M., Schindler, D.E., 2004. Climatic effects on the phenology of lake
processes. Glob. Change Biol. 10, 1844e1856.
Winograd, I.J., Landwehr, J.M., Coplen, T.B., Sharp, W.D., Riggs, A.C., Ludwig, K.R.,
Kolesar, P.T., 2006. Devils Hole, Nevada, d18O record extended to the midHolocene. Quat. Res. 66, 202e212.
Western Regional Climate Center (WRCC), 2012. Individual Station Data for Cherry
Valley Dam, CA, Station #041697. http://www.wrcc.dri.edu/.
Yuan, F., Linsley, B.K., Lund, S., McGeehin, J.P., 2004. A 1200 year record of hydrologic variability in the Sierra Nevada from sediments in Walker Lake. Nevada.
Geochem. Geophys. Geosys. 5, 1e13.
Zhang, Y., Wallace, J.M., Battisti, D.S., 1997. ENSO-like interdecadal variability: 190093. J. Climate 10, 1004e1020.