A Holocene Record of Endogenic Iron and Manganese 

Precipitation, Isotopic Composition of Endogenic 

Carbonate, and Vegetation History in a Lake-Fen Complex 

in Northwestern Minnesota

By Walter E. Dean and Lisa A. Doner

Open-File Report 2011–1106

U.S. Department of the Interior

U.S. Geological Survey
ii

U.S. Department of the Interior

KEN SALAZAR, Secretary

U.S. Geological Survey

Marcia K. McNutt, Director

U.S. Geological Survey, Reston, Virginia: 2011

For product and ordering information:

World Wide Web: http://www.usgs.gov/pubprod

Telephone: 1-888-ASK-USGS

For more information on the USGS—the Federal source for science about the Earth,

its natural and living resources, natural hazards, and the environment:

World Wide Web: http://www.usgs.gov

Telephone: 1-888-ASK-USGS

Suggested citation:

Dean, W.E., and Doner, L.A., 2011, A Holocene record of endogenic iron and manganese precipitation, isotopic 

composition of endogenic carbonate, and vegetation history in a lake-fen complex in northwestern Minnesota: 

U.S. Geological Survey Open-File Report 2011–1106, 41 p.

Any use of trade, product, or firm names is for descriptive purposes only and does not imply 

endorsement by the U.S. Government. 

Although this report is in the public domain, permission must be secured from the individual 

copyright owners to reproduce any copyrighted material contained within this report.
iii

Contents

Abstract ......................................................................................................................................................................... 1

Introduction.................................................................................................................................................................... 2

Methods......................................................................................................................................................................... 4

Results........................................................................................................................................................................... 7

Little Shingobee Lake ................................................................................................................................................ 7

Chronology............................................................................................................................................................. 7

Sediment Components........................................................................................................................................... 9

Stable Isotopes of Carbon and Oxygen.................................................................................................................17

Pollen Zones .........................................................................................................................................................21

Sediment Components in Little Shingobee Fen........................................................................................................23

Discussion ....................................................................................................................................................................28

Carbonate and Detrital Clastic Deposition ................................................................................................................28

Endogenic Fe and Mn...............................................................................................................................................29

Stable Isotopes .........................................................................................................................................................29

Pollen Zones.............................................................................................................................................................34

Conclusions..................................................................................................................................................................36

Acknowledgments ........................................................................................................................................................37

References Cited..........................................................................................................................................................37

Appendix 1. Concentrations of Major, Minor, and Trace Components in Core LSL–B .......................................... link

Appendix 2. X-Ray Diffraction Intensities and Percentages of Dominant Minerals in Core LSL–B ....................... link

Appendix 3.1. Results of Carbon and Oxygen Isotope Analyses of Ostracodes...................................................... link

Appendix 3.2. Results of Carbon and Oxygen Isotope Analyses of Pisidium........................................................... link

Appendix 4. Results of Carbon Analyses of Organic Carbon in Core LSL–B......................................................... link

Appendix 5. Percentages of Major Pollen Types in Core LSL–B .......................................................................... link

Appendix 6. Concentrations of Major, Minor, and Trace Components in Core LSF–10 ........................................ link

Figures

1. A, Modern lakes of the Shingobee River Headwaters Area. B, Inferred extent of glacial Lake Willobee

formed when a debris flow dammed the Shingobee River ...................................................................................... 3

2. Little Shingobee Lake and Little Shingobee Fen. Transects of cores of Carter and others (1997), and 

locations of cores LSL–B and LSF–10 also are shown ........................................................................................... 5

3. Age-depth model for Little Shingobee Lake core LSL–B from Little Shingobee Lake based on AMS 

radiocarbon dates ................................................................................................................................................... 9

4. Sediment components in Little Shingobee Lake and Little Shingobee Fen........................................................... 11

5. Major and minor components in sediments in core LSL–B from Little Shingobee Lake relative to depth

in core ................................................................................................................................................................... 12

6. Percentages of allogenic (detrital) and endogenic components in sediments in core LSL–B from Little 

Shingobee Lake relative to depth in core. Ages of sediments from the age model in figure 3 are shown at the 

right of the diagram. Pollen zones are for cores from Little Shingobee Lake (Doner, this report), Shingobee 

Lake (Locke, 1995), and Elk Lake (Whitlock and others, 1993) ............................................................................ 14

7. Percentages of total, detrital, and endogenic iron (Fe) and manganese (Mn) in sediments in core LSL–B 

from Little Shingobee Lake relative to depth in core. See text and figure 4 for methods of calculation of detrital 

and endogenic Fe and Mn. Ages of sediments from the age model in figure 3 are shown at the right of the Mn 

diagram................................................................................................................................................................. 16
iv

8. Stable isotope ratios of carbon and oxygen (δ13C and δ18O) in the carbonate shells of the fingernail clam, 

Pisidium, and ostracodes picked from sediments in core LSL–B from Little Shingobee Lake relative to depth in 

core....................................................................................................................................................................... 18

9. Stable isotope ratios of carbon (δ13C) in organic carbon and percent organic carbon in sediments in core 

LSL–B from Little Shingobee Lake relative to depth in core ................................................................................ 20

10. Pollen diagram for major pollen types in core LSL–B from Little Shingobee Lake relative to depth in core ........ 22

11. Major and minor components in sediments in core LSF–10 from Little Shingobee Fen relative to depth in 

core ..................................................................................................................................................................... 24

12. Percentages of allogenic (detrital) and endogenic components in sediments in core LSF–10 from Little

Shingobee Fen relative to depth in core.............................................................................................................. 25

13. Percentages of total, detrital, and endogenic iron (Fe) and manganese (Mn) in sediments in core LSF–10 

from Little Shingobee Fen relative to depth in core. See text and figure 4 for methods of calculation of detrital 

and endogenic Fe and Mn................................................................................................................................... 27

14. Cross plots of values of δ18O compared with δ13C in endogenic carbonate (marl) from Elk, Williams, and 

Shingobee Lakes, and in ostracodes from Little Shingobee Lake for three time slices: 10 cal ka, 4 cal ka, and 

2 cal ka. A fourth time slice (7 cal. ka) is added for Elk and Little Shingobee Lakes for which there is better time

control.................................................................................................................................................................. 31

Table

1. Radiocarbon dates from Little Shingobee Lake core LSL–B................................................................................... 8
v 

Abbreviations Used in This Report

cal. ka thousand calendar years old

cal. yr calendar year

cm centimeter

cmblf centimeters below lake floor

g gram

ka thousands of years old

ky thousands of years

mm millimeter

mm yr–1

 millimeters per year

AMS accelerator mass spectrometer

BP before present

DIC dissolved inorganic carbon

GMWL global meteoric water line

IC inorganic carbon

LMWL local meteoric water line

OC organic carbon

SHAEP Shingobee Headwaters Aquatic Ecosystem Project

TC total carbon

USGS U.S. Geological Survey

VPDB Vienna Pee Dee belemnite

XRD X-ray diffraction
1

A Holocene Record of Endogenic Iron and Manganese 

Precipitation, Isotopic Composition of Endogenic 

Carbonate, and Vegetation History in a Lake-Fen 

Complex in Northwestern Minnesota

By Walter E. Dean1 and Lisa A. Doner2

1

U.S. Geological Survey, MS 980 Federal Center, Denver, Colorado 80225

2

Center for the Environment, MSC6317, Plymouth State University, Plymouth, New Hampshire 03264

Abstract

Little Shingobee Lake and Fen are part of an extensive network of lakes and wetlands in the 

Shingobee River headwaters area of northwestern Minnesota. Prior to about 9800 radiocarbon years, 

most of the lakes in the Shingobee watershed area were interconnected to form glacial Lake Willobee.

From 9800 to 7700 radiocarbon years, the level of Lake Willobee fell as a result of breaching of a dam, 

leaving small separated basins containing the existing lakes and wetlands. 

The dominant components in the sediments in a 9-meter core from Little Shingobee Lake 

(LSL–B), and lacustrine sediments under 3.3 meters of peat in a 17-meter core from Little Shingobee 

Fen (LSF–10) are detrital clastic material, endogenic CaCO3, and organic matter. The detrital fraction in 

the Holocene section in core LSL–B varies considerably from 7 weight percent to 82 weight percent and 

closely parallels the concentration of detrital quartz measured by X-ray diffraction. The CaCO3

concentration, which also varies considerably from 10 weight percent to 70 weight percent, is generally 

antithetic to the detrital concentration owing to the dilution of detrital material by CaCO3, particularly 

during the early to middle Holocene (about 9000–6500 calendar years). The organic-matter content 

varies from 5 weight percent to 25 weight percent and, together with CaCO3, serves to dilute the 

allogenic detrital fraction.

In both cores almost all of the iron (Fe) and manganese (Mn) is in endogenic minerals, presumed 

to be oxyhydroxide minerals, that are important components throughout the core; little Fe and Mn are

contributed by detrital aluminosilicate minerals. The endogenic Fe mineral, calculated as Fe(OH)3, 

forms a larger percentage of the sediment than endogenic organic material throughout most of the 

Holocene section in the LSL–B core and in the lacustrine sediments below the peat in the LSF–10 core. 

Biogenic silica as opal (biopal; diatom debris) was not measured, but the average calculated biopal is 

5 percent in the LSL–B core and 15.5 percent in the LSF–10 core. 

Values of δ18O in mollusk (Pisidium) and ostracode shells increase by only about 20 per mil 

from the bottom to the top of the LSL–B core (about 12600–2200 calendar years). The remarkably 

constant oxygen-isotope composition throughout the Holocene suggests that environmental conditions 

affecting values of δ18O (temperature, salinity, composition of the water, composition of precipitation) 

did not change greatly. Values of δ13C in carbonate shells generally decreased by about 2 per mil 

from 9000 calendar years to 6000 calendar years, but they did not increase in organic carbon. This 
2

mid-Holocene increase in δ13C in shells but not in organic carbon is likely due to an increase in 

residence time.

A late Pleistocene forest dominated by spruce was replaced in the early Holocene by a pine 

forest. The pine forest migrated east during the middle Holocene and was replaced by an open 

sagebrush–oak savanna. The western migration of forests into northwestern Minnesota is marked first 

by a hardwood forest and finally a pine forest. 

Introduction

The lakes and wetlands of the Shingobee River headwaters area in northwestern Minnesota (fig. 

1A) have been intensively studied by the U. S. Geological Survey (USGS) and several universities for 

more than 20 years (Winter, 1997). The first initiative, the Interdisciplinary Research Initiative (IRI), 

focused mainly on Williams and Shingobee Lakes. Preliminary results of numerous IRI studies have 

been published in Winter (1997). Most of the studies concentrated on the waters of the lakes and 

watershed; less work has been done on the sediments. However, some results from analyses of surface 

sediments and longer sediment cores have been presented elsewhere (Schwalb and others, 1995; Dean 

and Bradbury, 1997; Locke and Schwalb, 1997; Dean and Schwalb, 2002; Schwalb and Dean, 2002). In 

recent years research expanded to include the entire Shingobee headwaters under the Shingobee 

Headwaters Aquatic Ecosystem Project (SHAEP). Research continued to be focused on waters and the 

watershed, but in addition long cores that collected the entire Holocene sediment sections were taken 

from Little Shingobee Lake, Little Shingobee Fen, and Steel Lake (Carter and others, 1997; Wright and 

others, 2004; Tian and others, 2005). Prior to about 9.8 thousand radiocarbon years ago (radiocarbon 

ka), most of the lakes in the Williams and Shingobee watershed area were interconnected to form glacial 

Lake Willobee (fig. 1B; Schwalb and Dean, 2002), which apparently formed when a debris flow 

dammed the Shingobee River (R.C. Melchior, 2010, unpub. data). From 9.8 to 7.7 radiocarbon ka, the 

level of glacial Lake Willobee fell, as a result of breaching of the dam, and left small separated basins 

containing the existing lakes and wetlands. At that time, Williams Lake became hydrologically isolated 

(it had no surface-water connection). The other lakes have surface inlets and outlets but also store a 

large component of groundwater from both shallow and deep sources (Rosenberry and others, 1997; 

Dean and others, 2003). 
3

Figure 1. A, Modern lakes of the Shingobee River Headwaters Area. B, Inferred extent of glacial Lake Willobee 

formed when a debris flow dammed the Shingobee River.

Williams, Shingobee, and Steel Lakes are three of the lakes occupying basins on the floor of 

glacial Lake Willobee. The Holocene sediments in these lakes contain abundant endogenic carbonate 

(marl) (Dean and Schwalb, 2002; Schwalb and Dean, 2002; Tian and others, 2005; Dean, 2010, unpub. 

data). For sediment description, magnetic susceptibility, loss-on-ignition, pollen stratigraphy, and 

ostracode assemblages in cores from Williams and Shingobee Lakes, see Schwalb and others (1995) and 

Locke and Schwalb (1997). The SHAEP area has numerous seeps that precipitate an orange ferric 

hydroxide floc, and the groundwaters, especially in a deep aquifer (several tens of meters to more than 

130 meters (m) below the surface), contain high concentrations of iron (Fe) and manganese (Mn) (Dean 

and others, 2003). The sediments of Shingobee Lake are unusual for a lake in a glaciated northtemperate region in that precipitation of Fe and Mn minerals, presumed to be X-ray amorphous 

oxyhydroxides, occurs at spring and fall overturn (Dean and others, 2003). As a result, the surface 

sediments of Shingobee Lake contain concentrations of Fe and Mn that are 7 and 27 times higher than 

can be explained by contributions of Fe and Mn from detrital aluminosilicates, and even higher 
4

concentrations occur at depth in cores (Dean and Schwalb, 2002). Concentrations of dissolved Fe and 

Mn in the hypolimnion of Shingobee Lake at the end of summer stratification are several hundred times 

higher than after spring and fall overturn, and they are much higher than in shallow groundwaters. 

However, groundwater from deep wells (as deep as 130 m) contains high concentrations of Fe and Mn 

(as much as 14 parts per million (ppm) Fe and 0.5 ppm Mn; that is, they are major ions), and this deep 

groundwater is the most likely source of high concentrations of Fe and Mn that accumulate in the 

hypolimnion of Shingobee Lake at the end of summer stratification (Dean and others, 2003). 

The Holocene dominant sediment components in Shingobee Lake are allogenic detrital material 

and endogenic CaCO3, organic matter, Fe and Mn oxyhydroxides, and an unknown amount of biogenic 

silica (diatom debris). Because Little Shingobee Lake is in the same groundwater regime as Shingobee 

Lake, we assumed that the same would be true of the Holocene sediments in Little Shingobee Lake and 

the lake sediments that underlie 3.3 m of peat in Little Shingobee Fen. The purposes of this report are 

(1) to document the composition of sediment components in Little Shingobee Lake and Fen, and 

particularly whether excess (endogenic) Fe and Mn occur in the sediments; (2) to document the 

vegetation and climatic history of the region on the basis of pollen; and (3) to determine the sources of 

organic matter and how these sources may have changed with changes in productivity, vegetation, and 

hydrology. The dominance of endogenic Fe and Mn as compared with detrital Fe and Mn in the 

sediments of Shingobee Lake is highly unusual, and we wanted to see if chemical conditions in Little 

Shingobee Lake would duplicate endogenic formation of Fe and Mn. The vegetation history of the 

Shingobee River headwaters area is based on studies of Shingobee and Williams Lakes, but the age 

model for these lakes is questionable. Isotopic composition of organic matter in Williams and 

Shingobee Lakes shows distinct changes with time. In order to document the timing of events and 

changes recorded in the sediments, a good age model was needed, and we attempted to provide it. 

Methods

Core LSL–B from Little Shingobee Lake was cored from the ice in 1998 with a modified squarerod Livingstone piston corer as an extension of a transect of cores across Little Shingobee Fen (Carter 

and others, 1997; transect B-B′ and D-D′ in fig. 2). The top meter of the core was too soupy to be 

recovered by the Livingstone corer, but the rest of the core was recovered; it consists of marl and gyttja. 

Core LSF–10 was cored from the surface of the fen along transect D-D′ of Carter and others (1997). The 

top 3.3 m of the core consists of peat, but the rest of the core consists of marl, bottoming in sand. The 

cores are stored at the Limnological Research Center, University of Minnesota. 
5

Figure 2. Little Shingobee Lake and Little Shingobee Fen. Transects of cores of Carter and others (1997), and 

locations of cores LSL–B and LSF–10 also are shown.

Concentrations of total carbon (TC) and inorganic carbon (IC) were determined by coulometric 

titration of CO2 following extraction from the sediment by combustion at 950°C and acid volatilization, 

respectively (Engleman and others, 1985). Weight percent IC was converted to weight percent CaCO3

by dividing by 0.12, the fraction of carbon in CaCO3. Organic carbon (OC) was determined as the 
6

difference between TC and IC. The accuracy and precision of both TC and IC, determined from 

hundreds of replicate standards, usually are better than 0.10 wt percent. 

Semiquantitative estimates of mineral contents in aliquots of carbon samples from core LSL–B 

were determined by standard X-ray diffraction (XRD) techniques (Moore and Reynolds, 1989). Each 

sample was packed into an aluminum holder and scanned from 15° to 50° 2Θ at 2° 2Θ/minute (min) by 

using Ni-filtered, Cu-Kα radiation at 45 kilovolts and 30 milliamps. For each sample, raw XRD peak 

intensities, in counts per second, of the main peaks of minerals detected in each sample were converted 

to semiquantitative percentages by dividing the main peak intensity of a mineral by the sum of the main 

peak intensities of all minerals. These percentage calculations should be used with caution because they 

do not reflect the different X-ray mass absorption characteristics of different minerals.

Samples were analyzed for 40 major, minor, and trace elements in a multiacid digest of sediment 

samples by inductively coupled argon-plasma, atomic emission spectrometry (ICP–AES) by SGS 

Minerals Services, Toronto, Canada. Rock standards (USGS) were included with the sediment samples, 

and 5 percent of the samples were analyzed in duplicate. The precision, determined by analyzing rock 

standards and duplicate sediment samples, was better than 10 percent, and usually better than 5 percent

at a concentration of 10 times the limit of detection. Only the results of analyses of aluminum (Al), 

titanium (Ti), iron (Fe), and manganese (Mn) will be discussed in this report. 

The bulk-sediment samples remaining after XRD and ICP analyses were decalcified with 

hydrochloric acid for measurement of carbon isotopes in organic carbon. Therefore, measurements of 

ratios of stable isotopes of carbon and oxygen were made only on picked mollusk (the fingernail clam 

Pisidium) and ostracode shells. No attempt was made to pick only one species of ostracode or to 

separate juveniles from adults. Consequently, the isotope analyses on ostracode valves are on mixed 

adults and juveniles and mixed species, most likely Candona acuta and Candona ohioensis, which are 

the most common ostracodes in Shingobee Lake sediments (Schwalb and Dean, 2002). Powdered 

samples were reacted with dehydrated phosphoric acid under vacuum at 70°C. The resulting CO2 was 

then dehydrated and purified in a high-vacuum gas-transfer system, and the isotope ratios determined 

with an isotope-ratio mass spectrometer. 

Stable carbon-isotope ratios of bulk organic matter were determined by standard techniques on 

decalcified samples (Dean and others, 1986). A powdered sample was reacted with buffered acetic acid 

for 24 hours to dissolve carbonate minerals. The residue was then centrifuged, decanted, washed three 

times, and dried. The residue was then combusted at 1,000°C with copper oxide in a sealed quartz tube. 

The resulting CO2 was then dehydrated and purified in a high-vacuum gas-transfer system, and the 

isotope ratios were determined with an isotope-ratio mass spectrometer. Duplicate analyses were made 

on all samples.

Results of carbon and oxygen analyses are reported in the usual per mil (‰) δ−notation relative 

to the Vienna Pee Dee Belemnite (VPDB) marine-carbonate standard for carbon and oxygen: 

δ‰=[(Rsample/RVPDB)−1]×103

,

where R is the ratio (13C: 12C) or (18O: 16O). Precision is ±0.1 per mil for oxygen and ±0.06 per mil for 

carbon in carbonate and on the basis of duplicate analyses, it is usually better than 0.2 per mil for 

organic carbon. 

Seventeen samples were analyzed for their contained pollen. Samples for pollen analysis were 

prepared by using standard procedures (Faegri and others, 1989). Spores of Lycopodium clavatum from 

Lund, Sweden, were added to each 0.5 gram (g) sample to provide a control basis for concentration 

estimates. Prepared pollen samples are stored with glycerin in capped glass vials. About 40–100 
7

terrestrial pollen grains were counted for each of the 17 levels, and pollen percentages were based on the 

sum of terrestrial pollen. 

Pollen identifications and counts were completed by using an Olympus BH2 microscope with a 

combined ocular and dry-lens magnification factor of 500. To avoid problems of differential sorting, 

pollen grains were identified along a minimum of five evenly spaced transects, covering either half the 

slide from midline to outer edge, or the full slide, until a minimum of 300 grains were identified, 

excluding spores, aquatic types, and pre-Quaternary types. The pollen types are grouped as trees, 

shrubs, herbs, aquatic, spores, pre-Quaternary, and indeterminate. Pollen percentages for each group 

were calculated relative to a total sum of trees, shrubs, herbs, and indeterminate types. Total pollen 

concentrations are derived from grains per gram of dry sediment and include all vascular pollen and 

spore types. Cyperaceae (sedges) in the temperate regions occur commonly as aquatic species and so are 

grouped as aquatic pollen. The multivariate analysis program CONISS (Grimm, 1987) generated 

Euclidean-distance dissimilarity values of terrestrial pollen and spores with more than 1 percent

maximum occurrence (Overpeck and others, 1985). These values are displayed as hierarchical, 

polythetic agglomerative dendrograms with minimum-variance clustering so that samples that are most 

alike (least dissimilar) form tight clusters close to the axis (Gaugh, 1982). The dendogram defines the 

pollen zones. 

The pollen zonations defined for Little Shingobee Lake resemble those of Whitlock and others 

(1993) for Elk Lake just north of the SHAEP area in the headwaters of the Mississippi River (fig. 1A, 

inset), and Williams and Shingobee Lakes (Locke, 1995; Locke and Schwalb, 1997), but the 

chronologies of these zones are slightly different (discussed below). 

Results

Little Shingobee Lake

Chronology

A total of 31 accelerator mass spectrometer (AMS) radiocarbon dates were obtained from

samples of wood, mollusks, and bulk sediment from Little Shingobee Lake (table 1). Targets for AMS 

analyses were prepared in the USGS Radiocarbon Laboratory, Reston, Va., (WW numbers in table 1) 

and analyzed at the Lawrence Livermore Center for Accelerator Mass Spectrometry (CAMS). Ages 

were calibrated by using both the University of Washington’s radiocarbon calibration program CALIB 

4.3 (Stuiver and others, 1998) and the University of Cologne’s online radiocarbon calibration program 

CalPal (http://www.calpal.de) (table 1). A depth-age diagram (fig. 3) shows that dates from mollusks are 

generally older than those from wood or bulk sediment, particularly at the base of the core, owing to the 

hard-water carbonate reservoir effect. For five horizons where there are dates from both mollusks and 

wood or sediment, the maximum difference is 1,130 years and the average difference is 338 years. To 

determine depth-age relations, we used the regression shown in figure 3 for wood and sediment only. 

All ages are expressed in thousands of calendar years before present (cal. ka BP), where “present” is AD 

1950 and ranges of ages are expressed in thousands of years (k.y.).
8

Table 1. Radiocarbon dates from Little Shingobee Lake core LSL–B.

[AMS, accelerator mass spectrometer; cmblf, centimenters below lake floor; cal. ka BP, thousands of calendar years before 

present; sed., sediment WW#, U.S. Geological Survey Radiocarbon Laboratory number; yr BP, years before present]

Depth Material Dated on USGS Radiocarbon Age AMS 14C Error CALIB 

Age

CalPal 

Age 

(cmblf) Lab Numbers 

(WW#) (yr BP) (years) (cal. ka 

BP)

(cal. ka 

BP)

419.5 wood 10/4/1998 1916 4655 60 5.42 5.40

658-661.5 wood 10/4/1998 1917 10810 95 12.89 12.91

810-812 cone 10/4/1998 1918 10175 80 11.82 11.83

840-842 wood 10/4/1998 1919 10300 140 12.25 12.11

845.5-846 wood 10/4/1998 1920 10305 70 12.24 12.17

871-872 wood 10/4/1998 1921 10510 120 12.47 12.40

915-917 wood 10/4/1998 1922 10820 100 12.90 12.81

420-422 bulk sed. 12/29/1998 2104 5400 50 6.25 6.20

659-661 bulk sed. 12/29/1998 2103 10710 50 12.85 12.90

915-917 bulk sed. 12/29/1998 2101 12840 50 15.48 15.37

810-812 bulk sed. 12/29/1998 2102 16180 100 19.30 19.30

840-842 bulk sed. 12/29/1998 2109 13000 40 15.63 15.85

120-122 bulk sed. 6/19/1999 2359 2915 55 3.08 3.07

170-172 bulk sed. 6/19/1999 2360 3170 55 3.38 3.40

230-232 bulk sed. 6/19/1999 2361 3395 55 3.64 3.65

380-382 bulk sed. 6/19/1999 2362 5605 85 6.37 6.41

430-432 bulk sed. 6/19/1999 2363 5505 95 6.29 6.30

542-544 bulk sed. 6/19/1999 2364 6555 70 7.43 7.47

602-604 bulk sed. 6/19/1999 2365 8125 70 9.03 9.10

619-621 bulk sed. 6/19/1999 2366 8440 70 9.47 9.44

770-772 bulk sed. 6/19/1999 2367 16230 110 19.36 19.40

409-411 mollusks 2/16/1999 2150 4970 40 5.68 5.70

840-842 mollusks 2/16/1999 2151 11350 50 13.25 13.24

870-872 mollusks 2/16/1999 2152 12250 40 14.21 14.32

120-122 mollusks 7/13/1999 2374 2410 50 2.36 2.37

173 mollusks 7/13/1999 2375 3270 50 3.47 3.51

380-382 mollusks 7/13/1999 2376 4940 40 5.66 5.67

427.5 mollusks 7/13/1999 2377 5660 50 6.42 6.45

542-544 mollusks 7/13/1999 2378 7000 50 7.81 7.84

602-604 mollusks 7/13/1999 2379 8930 50 10.15 10.06

619-621 mollusks 7/13/1999 2380 8730 50 9.70 9.72
9

Figure 3. Age-depth model for Little Shingobee Lake core LSL–B from Little Shingobee Lake based on AMS 

radiocarbon dates.

Sediment Components

The dominant components in the sediments of Little Shingobee Lake are detrital clastic material 

(aluminosilicates), endogenic CaCO3, biogenic silica (biopal), and organic matter (fig. 4; appendix 1). A 

record of clastic influx into the lake was provided by measured concentrations of aluminum (Al) and 

titanium (Ti) (fig. 5) because their concentrations are largely unaffected by weathering and 

postdepositional alteration. The quartz content, determined by X-ray diffraction, provides another, 

independent measure of the detrital fraction (fig. 5). X-ray diffraction analyses of the high-detrital 

sections of the core show that they consist predominantly of quartz with minor dolomite, feldspar, and 

calcite, in that order of abundance (appendix 2). The percentage of dolomite in the high-detrital sections 
10

is always higher than the percentages of feldspar and calcite, suggesting that the dolomite is detrital, 

derived from Paleozoic carbonate rocks in the drainage basin. Some of the calcite in these sections also 

may be detrital, but most is probably endogenic as in the low-detrital sections. A record of endogenic 

CaCO3 is calculated from measured concentrations of IC and XRD calcite. A record of endogenic 

organic matter (OM) is provided by measured concentrations of OC. Biogenic silica calculated as opal 

(biopal) was calculated as the residual from the sum of all other components (fig. 4; appendix 1; 

discussed below). 
11

Figure 4. Sediment components in Little Shingobee Lake and Little Shingobee Fen.
12

Figure 5. Major and minor components in sediments in core LSL–B from Little Shingobee Lake relative to depth in core.
13

The total detrital clastic fraction was calculated by assuming that the detrital rock debris in Little 

Shingobee Lake sediments has the same Al concentration as the carbonate- and organic-free sediment 

entering Little Shingobee Lake from the inlet at the south end of the lake (fig. 2) (6.8 percent Al). The 

calculated detrital fraction in core LSL–B varies considerably from 7 percent to 82 percent and closely 

parallels the concentration of detrital quartz measured by XRD (fig. 5). The CaCO3 concentration, 

calculated from the measured IC concentration, also varies considerably from 10 percent to 70 percent 

and is generally antithetic to the detrital concentration (fig. 5). The CaCO3 maxima between 660 and 

420 centimeters (cm) depth (about 9–6.5 cal. ka) and 250 to 180 cm (about 4–3 cal. ka) are most likely 

due to dilution of clastic material with CaCO3 rather than a decrease in the influx of detrital material.

The OM content, equivalent to loss on ignition at 550°C, is twice the OC content, a relation that has 

been demonstrated for many Minnesota lakes (Dean, 1974) and, in particular, in cores from Shingobee 

and Williams Lakes (Dean and Schwalb, 2002). The OM content in the core LSL–B varies from 

5 percent to 25 percent (average, 6.6 percent) and, together with CaCO3, serves to dilute the allogenic 

detrital fraction (fig. 5). The detrital Fe and Mn concentrations, expressed as weight percent Fe2O3 and 

MnO2 (fig. 5), were calculated by assuming that the detrital fraction in Little Shingobee sediments has 

the same Fe/Al and Mn/Al ratios as carbonate- and organic-free inlet stream inlet sediments (0.31 and 

0.003, respectively; fig. 4). 

In the top 50 cm of the core LSL–B (100–150 cm) the sum of the major components of allogenic 

detrital clastic material, endogenic CaCO3, and OM totals about 94 percent, leaving about 6 percent of 

other endogenic material (fig. 6). Detrital clastic material and CaCO3 clearly dominate the sediments in 

Little Shingobee Lake. In most of the core, detrital material dominated (60–70 percent) but percent

CaCO3 had two distinct pulses, one during the early Holocene (700–450 cm; about 10–7 cal ka; fig. 6), 

and a second during the late Holocene (250–150 cm; about 4–3 cal. ka). The percentage of OM is 

relatively minor throughout most of the section except for a pulse of up to 10 percent coincident with 

the late Holocene pulse in CaCO3 (fig. 6).
14

Figure 6. Percentages of allogenic (detrital) and endogenic components in sediments in core LSL–B from Little 

Shingobee Lake relative to depth in core. Ages of sediments from the age model in figure 3 are shown at the 

right of the diagram. Pollen zones are for cores from Little Shingobee Lake (Doner, this report), Shingobee 

Lake (Locke, 1995), and Elk Lake (Whitlock and others, 1993).
15

Because unusually high concentrations of endogenic Fe and Mn oxyhydroxides occur in the 

sediments of Shingobee Lake (Dean and Schwalb, 2002; Dean and others, 2003), they likely form a 

substantial part of the 6 percent of other endogenic components in Little Shingobee Lake. We define 

“excess” concentrations of Fe and Mn as the difference between the measured total Fe and Mn 

concentrations and the calculated detrital Fe and Mn concentrations (fig. 4). We then assume that the 

excess Fe and Mn are present as endogenic minerals, most likely as oxyhydroxides. The most oxidized 

species of Fe and Mn in natural waters are Fe(OH)3 and Mn(OH)2 (Wetzel, 2001), and we calculated 

excess Fe and Mn as these species for mass balance purposes. On the basis of these calculations, the 

average concentrations of excess Fe and Mn are as much as 2.6 and 10 times higher than can be 

explained by average contributions of Fe and Mn from detrital aluminosilicates and can be much higher 

(fig. 7). In fact, almost all of the Mn is endogenic (fig. 7). Endogenic Fe(OH)3 forms a larger percentage 

of the sediment than endogenic OM throughout most of the section (fig. 6). The section between 250 

and 150 cm depth (about 4–3 cal. ka) is dominated by endogenic components CaCO3, Fe(OH)3, OM, 

and biopal (discussed below) (fig. 6).
16

Figure 7. Percentages of total, detrital, and endogenic iron (Fe) and manganese (Mn) in sediments in core LSL–B 

from Little Shingobee Lake relative to depth in core. See text and figure 4 for methods of calculation of detrital 

and endogenic Fe and Mn. Ages of sediments from the age model in figure 3 are shown at the right of the Mn 

diagram. 
17

Diatoms are a major component of the phytoplankton in both Williams and Shingobee Lakes 

(LaBaugh, 1997) and probably also in Little Shingobee Lake. Therefore, biogenic silica as opal (biopal) 

is likely a major part of the other endogenic material, but we do not have direct measurements of biopal. 

We calculated weight percent biopal as the difference between 100 percent and the sum of percentage

detrital + percentage CaCO3 + percentage OM + percentage Fe(OH)3 + percentage Mn(OH)2 (fig. 4). The 

average percentage of biopal is 5.1 percent but the total percentage is as high as 15.5 percent (fig. 6). 

Stable Isotopes of Carbon and Oxygen

Values of δ18O and δ13C in Pisidium and ostracodes in Little Shingobee Lake show parallel 

trends but with considerable variation (fig. 8; appendix 3). Some of the variation in isotopic values in 

ostracodes is likely due to analysis of mixed adults and juveniles and mixed species. Values of δ18O 

increase only slightly from −9 to −8 per mil between 900 and 800 cm depth (about 12–11 cal. ka) to 

around −7 per mil at 100 cm, but with a large negative excursion to less than −10 per mil in values in 

Pisidium between 100 and 150 cm. This negative excursion is defined by four data points and must be 

real, but because Pisidium and ostracodes are benthic, we see no reason for such a large excursion only 

in Pisidium. Values of δ13C in ostracodes in pre-Holocene sediments (>700 cm) are −7 to −6 per mil 

before decreasing abruptly to −9.5 per mil at 650 cm. Values of δ13C show a broad, well-defined, early 

Holocene peak in both ostracodes and Pisidium, increasing from around −9 per mil at 650 cm (about 

9.5 cal. ka) to −8 between 600 and 400 cm (about 8.7 to 6.1 cal. ka). This 13C peak is approximately 

coincident with that of percent CaCO3 (fig. 5) but the 13C peak begins and ends later, suggesting that 

carbonate saturation occurs before isotopic equilibrium of dissolved inorganic carbon (DIC) is achieved. 

Values of δ13C in both ostracodes and Pisidium gradually decrease to about −9.5 per mil at 350 cm 

(about 5.5 cal. ka). The positive excursion in δ13C in Pisidium is based on only one data point and may 

be analytical error. 
18

Figure 8. Stable isotope ratios of carbon and oxygen (δ13C and δ18O) in the carbonate shells of the fingernail 

clam, Pisidium, and ostracodes picked from sediments in core LSL–B from Little Shingobee Lake relative to 

depth in core.
19

Values of δ13C in organic carbon in pre-Holocene sediments in Little Shingobee Lake (940–700 

cm depth, about 13–10 cal. ka) vary considerably, ranging from −32 per mil to −25 per mil (fig. 9; 

appendix 4). However, in Holocene sediments (<700 cm) variations in values of δ13C are much smaller 

(−30 per mil to −28 per mil), with a slight decrease of about 1.0 per mil upward in the section. The 

isotopic composition of OC does not correspond to any of the considerable variations in amount OC, 

especially in upper Holocene sediments (2 percent to >10 percent; fig. 9). In surface sediments in 

Minnesota lakes, those judged from plant pigments to be the most algal rich by Gorham and Sanger 

(1975) have values of δ13C of −30 per mil to −31 per mil (Dean and Stuiver, 1993). Those judged by 

pigments to be most oxidized have values of δ13C in the low −20s per mil. The most algal-rich surface 

sediments, and most 13C depleted, are also the most hydrogen rich (lipid rich) as judged by Rock-Eval 

pyrolysis (Dean and Stuiver, 1993). On the basis of these criteria, the organic matter in Little Shingobee 

Holocene sediments with values of δ13C of around −30 per mil should be algal. However, the increase in 

OC between 250 and 150 cm may be due to a greater influx of terrestrial organic matter, and the 

increase in δ13C from −29 per mil to −27 per mil may reflect that origin.
20

Figure 9. Stable isotope ratios of carbon (δ13C) in organic carbon and percent organic carbon in sediments in 

core LSL–B from Little Shingobee Lake relative to depth in core.
21

Pollen Zones

The basal sediments of late Pleistocene age (fig. 6; appendix 5) are dominated by spruce (Picea) 

and Cupressaceae (cypress). When the observed pollen types in core LSL–B are combined with those of 

Locke (1995) and Whitlock and others (1993), other abundant pollen types include Larix (larch), 

Gramineae (grass), Cyperaceae (sedge), and Artemisia (sagebrush). 

The Jack/Red Pine zone of early Holocene age (fig. 6) is dominated by Pinus banksiana and P. 

resinosa (fig. 10). The boundary between the Pine and Prairie zones was dated at 7680 radiocarbon 

years (6560 calendar years (cal. yr)) in a piece of wood from the Williams Lake core. Dates on all other 

pollen-zone boundaries reported by Locke (1995) (fig. 6) for Williams and Shingobee Lakes are based 

on other pollen stratigraphies from northwestern Minnesota. Pollen zone boundaries for Elk Lake 

reported by Whitlock and others (1993) were dated by varve counts and radiocarbon dates. 
22

Figure 10. Pollen diagram for major pollen types in core LSL–B from Little Shingobee Lake relative to depth in core.
23

The Prairie zone pollen taxa are dominated by Quercus (oak), grass, and sagebrush (fig. 10). 

During this mid-Holocene period, the forest-prairie border in northwestern Minnesota migrated at least 

100 km east of its present position, changing from pine forest to sagebrush-and-grass prairie and oak 

savanna (Whitlock and others, 1993). We use Wright’s (1976) term "prairie period" for the midHolocene dry episode in northwestern Minnesota, marked simply by dry conditions with no necessary 

connotation of temperature (discussed later). 

The return of the forest is marked by an increase in arboreal hardwood pollen taxa dominated by 

Quercus and Betula (birch) (fig. 10) but with important contributions from Ostrya (hornbeam) and pine. 

The hardwood forest was replaced by the return of a pine forest to northwestern Minnesota, this time 

dominated by white pine (P. strobus) but with substantial contributions from P. banksiana and P. 

resinosa. 

Sediment Components in Little Shingobee Fen

The components in the sediments of Little Shingobee Fen (core LSF–10) can be calculated by 

using the same assumptions as for sediments in Little Shingobee Lake (fig. 4; appendix 6), again 

assuming that the detrital clastic fraction in the fen has the composition as the carbonate- and organicfree sediments entering the south end of the lake (fig. 2). The calculated total detrital fraction of 

sediments in core LSF–10 is considerably lower than in core LSL–B throughout most of the Holocene 

section of the core below the top 330 cm of peat and above 1,400 cm depth, averaging 4.4 percent

(range, 1.3 percent to 31.0 percent) (fig. 11). CaCO3 is by far the dominant component with an average 

of 61 percent. Between 1,400 and 330 cm, the next most abundant components are biopal (average, 

16 percent) and OM (average, 13 percent) (fig. 12). Therefore, unlike the Holocene sediment section in 

core LSL–B with a dominant detrital component throughout most of its history, the sediment section in 

core LSF–10 was dominated by endogenic components (approximate ages to be determined below).
24

Figure 11. Major and minor components in sediments in core LSF–10 from Little Shingobee Fen relative to depth in core.
25

Figure 12. Percentages of allogenic (detrital) and endogenic components in sediments in core LSF–10 from Little 

Shingobee Fen relative to depth in core.
26

Because the detrital fraction in core LSF–10 is so low, the contributions of detrital Fe and Mn 

are also low (expressed as Fe2O3 and MnO2; fig. 11). However, as in core LSL–B, almost all of the Fe 

and Mn present in the sediments is endogenic (expressed as Fe(OH)3 and Mn(OH)2; fig. 13), even in the 

high-detrital clastic part of the section below 1,400 cm depth. On the basis of calculations of 

percentages of excess Fe and Mn (fig. 4), the average concentrations of excess Fe and Mn in core 

LSF–10 are 114 and 193 times higher than can be explained by contributions of Fe and Mn from 

average detrital aluminosilicates (fig. 13). 
27

Figure 13. Percentages of total, detrital, and endogenic iron (Fe) and manganese (Mn) in sediments in core 

LSF–10 from Little Shingobee Fen relative to depth in core. See text and figure 4 for methods of calculation 

of detrital and endogenic Fe and Mn.
28

Discussion

Carbonate and Detrital Clastic Deposition

All present-day lakes and wetlands of the Shingobee River headwaters area were part of larger 

glacial Lake Willobee (fig. 1B). Once Lake Willobee drained, a sill separated Little Shingobee Lake and 

the lake that occupied the basin under the present-day fen. Little Shingobee Fen-lake filled with 11 m of 

mostly carbonate sediment (marl), culminating in 3.3 m of peat (fig. 12). The deepest part of the Little 

Shingobee Fen basin was 17 m, the thickness of the sediment in core LSF–10 (Carter and others, 1997). 

Wood from the sandy base of core LSF–9 (8 m) just east of LSF–10 was radiocarbon dated at 10345 

years (about 12300 cal. yr; Carter and others, 1997), which may be the age of the base of LSF–9, or the 

wood could be old and redeposited. If the wood was not redeposited, and 12.3 cal. ka is a true date for 

the base of LFS–9, then that would also be the date of the sandy base of core LSF–10, which is in good 

agreement with our radiocarbon model for the base of core LSL–B (fig. 3). The bottom of the Little 

Shingobee Fen lake basin along transect C-C′ (fig. 2) is mostly at about 6 m, and the bottom along 

transect B-B′ is mostly at about 9 m (Carter and others, 1997). 

Assuming that the base of core LSF–10 is indeed the same as the wood-dated base of core 

LSF–9 (12.3 cal. ka), and the top of the basal clastic layer in LSF–10 (about 1,450 cm depth, fig. 12) is 

the same age as the top of the basal clastic layer in LSL–B (about 10 cal. ka at 7 m, fig. 6), then the 

average Holocene sedimentation rate in LSF–10 is 1.5 mm yr−1

, which is considerably faster than in 

either LSL–B (0.7 mm yr−1) or Shingobee Lake (0.9 mm yr−1; Schwalb and Dean, 2002). The basal 

detrital clastic layer in LSF–10 differs from that in LSL–B in that a 1-m-thick carbonate layer is 

sandwiched in the detrital layer in LSF–10 (fig. 12). That carbonate material (calcite and dolomite) may 

be detrital as is the minor CaCO3 (~20 percent) in the basal detrital layer in LSL–B (fig. 5), derived 

from calcareous glacial drift in the drainage basin (Itasca Moraine; Moores and Norton, 1997). The 

basal detrital layer in LSL–B has a higher percentage of dolomite (average, 16 percent) than calcite 

(average, 10 percent), further evidence of Paleozoic carbonate-derived calcareous drift.

Once glacial Lake Willobee drained, the Holocene sediments in Little Shingobee Fen-lake and 

Shingobee Lake (Schwalb and Dean, 2002) remained at a similar composition, consisting mostly of 

70–80 percent CaCO3, presumed to be mostly endogenic, with 330 cm of peat at the top of LSF–10 

(fig. 12). The sediments in core LSL–B are different in that the Holocene marl is interrupted by a 

2-m-thick detrital clastic layer (fig. 6), whereas detrital material in LSF–10 is very minor throughout the 

Holocene section (fig. 12). This difference may be because the sill between Little Shingobee Lake and 

Little Shingobee Fen-lake prevented detrital material from entering the fen-lake basin. Holocene 

sedimentation in Williams Lake was initially similar to that in the other three lakes, with 80 percent

CaCO3 in sediments above the detrital base (Dean and Schwalb, 2002). This early pulse of CaCO3

deposition may have been the result of postglacial leaching of CaCO3 from the calcareous Itasca

moraine (Moores and Norton, 1993). Unlike the other three lakes, however, Williams Lake became 

progressively closed hydrologically. As productivity in Williams Lake increased, decomposition of 

produced organic matter caused the hypolimnion to become more acidic, dissolving CaCO3 that was 

produced in the epilimnion (Dean and Schwalb, 2002). By about 5000 14C yr, the sedimentary organic 

carbon content in Williams Lake was greater than 10 percent (greater than 20 percent OM), and 

sedimentary CaCO3 was completely dissolved in the water column and surface sediments, as it is today. 
29

Endogenic Fe and Mn

As discussed above, the late Pleistocene influx of detrital material in Little Shingobee Lake 

(below 700 cm depth in core LSL–B; >10 cal. ka; figs. 5 and 6) is likely equivalent to the pulse of 

detrital material at the base of core LSF–10 (below 1,500 cm; figs. 11 and 12), with the difference that 

in LSF–10 there is a 1-m-thick CaCO3 layer between 1,550 and 1,650 cm (fig. 11). Therefore, it is likely 

that the peak in endogenic Mn in core LSF–10 at 1,500 cm (fig. 13) is equivalent to the endogenic Mn 

in core LSL–B at 660 cm (about 9.5 cal. ka; fig. 7), with similar bulk-sediment Mn contents (about 0.8 

percent). However, the endogenic Fe peak in core LSF–10 (up to 14 percent between 1,300 cm and 

1,500 cm; fig. 12) is much higher than in LSL–B (about 5 percent; fig. 7) and occurs after the endogenic 

Mn peak. As discussed above, prior to 9.8 radiocarbon ka (9.3 cal. ka), all of the lakes in the Shingobee 

River headwaters were connected as part of a larger glacial Lake Willobee (fig. 1B; Schwalb and Dean, 

2002). At that time the sediments of Williams, Shingobee, and Little Shingobee Lakes, and what is now 

Little Shingobee Fen, all contained high concentrations of Fe and Mn (Dean and Schwalb, 2002). 

The base of the varved sediment section in Elk Lake (10–8 varve ka) (fig. 1A inset), also 

contained high concentrations of Fe (about 5 percent) and Mn (about 2 percent). Dean (1993) suggested 

that the high concentrations of Fe and Mn in lower Holocene sediments in Elk Lake were caused by 

leaching of Fe and Mn from glacial till in the drainage basin early in the lake’s history. The most likely 

source of high concentrations of Fe and Mn that accumulate in the hypolimnion of Shingobee Lake 

today is not shallow groundwater but rather a deep till aquifer that overlies Cretaceous bedrock (Dean 

and others, 2003). Therefore, it seems likely that the source of Fe and Mn that precipitated as endogenic 

oxyhydroxide minerals that accumulated in the pre-Holocene sediments of Lake Willobee, represented 

by the sediments in Williams, Shingobee, and Little Shingobee Lakes and Little Shingobee Fen, was 

postglacial activation of that deep till aquifer. 

In observations of individual annual laminations (varves) from Elk Lake, an orange ferric 

hydroxide (?) gel usually forms distinct seasonal (summer) laminations (Anderson, 1993). The orange 

ferric hydroxide laminae are usually underlain by thinner, black laminae, probably composed of 

manganese oxyhydroxide. The common appearance of a fairly thick seasonal lamina of ferric hydroxide 

gel in Elk Lake varves indicates that much of the precipitated endogenic iron hydroxide is incorporated 

into the reduced sediments and not remineralized; the same is true of manganese oxyhydroxides. Ferric 

hydroxide colloids tend to be fairly stable (Tipping and others, 1981; Davison and Seed, 1983; Stauffer 

and Armstrong, 1986) which may explain why they survive in reducing hypolimnion and sediment pore 

waters. As in Elk Lake, the high concentrations of endogenic Fe and Mn in the sediments of the SHAEP 

lakes probably form at fall overturn when high concentrations of dissolved Fe and Mn that accumulate 

in the anoxic hypolimnion during summer stratification mix with oxygenated epilimnetic waters 

(Megard and others, 1993; Dean and others, 2003). 

Stable Isotopes

Once glacial Lake Willobee drained, the carbon and oxygen isotope composition of endogenic 

and ostracode-shell carbonate should be similar in lower Holocene sediments in all of the residual lakes 

on the floor of glacial Lake Willobee, at least initially. Apparently they were not (fig. 14). These 

differences are due mainly to vital and environmental effects. Benthic ostracodes reflect the isotopic 

composition of bottom waters. During summer stratification, hypolimnion δ18O in Shingobee Lake is 

about 1.5 per mil lower than the epilimnion δ18O (Kendall and others 1997), and the hypolimnion δ13C 

is about 3 per mil lower than the epilimnion DIC (R.G. Striegl, USGS, written commun., 2008). Culture 

experiments and studies of living ostracodes and modern lake waters have shown that some ostracode 
30

species calcify out of isotopic equilibrium with the water (Xia and others, 1997a; von Grafenstein and 

others, 1999). Also, if ostracodes molt throughout a wide range of temperatures (in different seasons), 

the isotopic signature reflects an integrated seasonal signal (Xia and others, 1997b). Oxygen-isotope 

values may display positive offsets from equilibrium of about 2 per mil. 
31

Figure 14. Cross plots of values of δ18O compared with δ13C in endogenic carbonate (marl) from Elk, Williams, and Shingobee Lakes, and in 

ostracodes from Little Shingobee Lake for three time slices: 10 cal ka, 4 cal ka, and 2 cal ka. A fourth time slice (7 cal. ka) is added for Elk and 

Little Shingobee Lakes for which there is better time control.
32

Ostracodes that grew in both Shingobee and Williams Lakes at the beginning of the Holocene 

have lower values of δ13C and higher values of δ18O than the marl (carbonate in bulk sediment) in which 

they occur (Schwalb and Dean, 2002), reflecting that positive equilibrium. Schwalb and Dean (2002) 

also found that all Holocene sediments in Shingobee Lake had values of δ18O in marl that were between 

3 per mil and 6 per mil lower than values in ostracodes, and in Williams Lake, marl values were 4 per 

mil to 9 per mil lower than in ostracodes. In fact, values of δ18O in Shingobee Lake marl changed very 

little throughout the Holocene section (about −10 per mil; fig. 14). Values of both δ18O and δ13C in 

ostracodes in Williams and Shingobee Lakes are much more variable than the marl values because 

several ostracodes species had to be used, and juvenile ostracodes were analyzed at levels where adults 

were not present (Schwalb and Dean, 2002). Also, if ostracodes molt throughout a wide range of 

temperatures (in different seasons), the isotopic signature reflects an integrated seasonal signal (Xia and 

others, 1997b). Ostracodes are present throughout the entire section in Shingobee Lake, but in Williams 

Lake ostracodes are present only in lower Holocene sediments, whereas bulk-sediment CaCO3 persisted 

until about 4 radiocarbon ka (4.5 cal. ka). For this study, only picked ostracodes and Pisidium were 

analyzed, and no attempt was made to separate ostracode species for analysis. On the basis of all of 

these considerations, the isotopic values for Little Shingobee marl deposited at 10 cal. ka (not measured) 

should have lower values of δ18O and higher values of δ13C than the ostracodes. 

The most obvious changes in postglacial Lake Willobee isotopic compositions are the marked 

increases in values of both δ18O and δ13C in Williams Lake marl between 10 cal. ka and 4 cal. ka (fig. 

14). This increase reflects the evolution of Williams Lake to a closed-basin lake. Today, a diagram of 

δ18O and δD of rain water in the Shingobee River headwaters area defines a local meteoric water line 

(LMWL) that is almost identical to the global meteoric water line (GMWL) (Kendall and others, 1997). 

Most river, lake, and groundwater samples, on the other hand, have values of δ18O that are higher than 

the LMWL; those higher values indicate a strong evaporation signal with Williams Lake, which has the 

highest values of δ18O. Values of δ13C and δ18O measured in samples collected during late summer on a 

transect along the Shingobee River and through Little Shingobee and Shingobee Lakes increase 

downstream along the Shingobee River transect from −11 per mil to −7 per mil for both δ13C and δ18O, 

indicating an evaporation trend and longer residence time downstream (Kendall and others, 1997). 

Williams Lake, on the other hand, had values of δ13C and δ18O that are enriched in 13C and 18O by about 

5 per mil and 6 per mil, respectively, relative to the most 13C- and 18O-enriched samples from Little 

Shingobee and Shingobee Lakes.

The remarkably constant oxygen-isotope composition throughout the Holocene in Shingobee 

Lake marl (Dean and Schwalb, 2002; Schwalb and Dean, 2002; fig. 14) and Little Shingobee Lake 

ostracodes (figs. 8 and 14) suggests that environmental conditions affecting values of δ18O (temperature, 

salinity, composition of the water) did not change greatly. The oxygen-isotope composition of Williams 

Lake marl did change markedly as the lake evolved from an open to a closed basin (fig. 14). 

In Shingobee Lake marl, there is a considerable increase in values of δ13C during the prairie 

period between 10 cal. ka and 4 cal. ka (about 4 per mil), then about a 2 per mil decrease between 4 cal. 

ka and 2 cal. ka (fig. 14). The mid-Holocene increase in values of δ13C (about 3 per mil) in Shingobee 

Lake is fairly broad and extends from about 7 radiocarbon ka (about 7.8 cal. ka) to 3.5 radiocarbon ka 

(about 3.8 cal ka) (Schwalb and Dean, 2002). The mid-Holocene increase in values of δ13C (about 2 per 

mil) in ostracodes in Little Shingobee Lake is also broad, increasing from about −9 per mil to −7 per 

mil, extending from about 9.5 cal. ka to about 5 cal. ka with a maximum of −7 per mil at about 7 cal. ka 

(figs. 8 and 14). By 4 cal. ka values of δ13C are back down to where they were at 9.5 ka (about −9 to 

−10 per mil; figs. 8 and 14). The marl in the varved sediments of Elk Lake also has a mid-Holocene 
33

maximum in values of δ13C that is even pronounced, going from −6 per mil at 10 cal. ka to +1 per mil at 

7 cal. ka (fig. 14). Values of both δ13C and δ18O are generally higher in marl from Elk Lake than in marl 

from Shingobee Lake due to the longer residence time of this 30-m-deep lake. 

The most commonly cited model for enrichment of surface waters by 13C is increased 

productivity with removal of 13C-depleted OC leaving reservoirs of 13C-enriched surface-water DIC and 

CO2 (Hollander and McKenzie, 1991; McKenzie and Hollander, 1993; Bernasconi and others, 1997; 

Hollander and Smith, 2001; Meyers and Teranes, 2001; Myrbo and Shapley, 2006). Today values of 

δ13C in surface waters of Shingobee and Little Shingobee Lakes are −11 to −12 per mil (R.G. Striegl, 

USGS, written commun., 2008). Values for waters in Williams Lake are much higher, about −4 per mil

for δ13C (R.G. Striegl, USGS, written commun., 2008). During the spring and summer (April to 

September), values of δ13C in DIC in the epilimnion increase by 1.40 per mil in Williams Lake and 1.24

per mil in Shingobee Lake because of photosynthetic uptake of lake CO2 that is enriched in 12C (Striegl 

and Michmerhuizen, 1998; Striegl and others, 2001). By late August of 1997, values of δ13C in CO2 in 

the epilimnion of Williams Lake were –10 per mil and those in the hypolimnion were −12 per mil; 

epilimnion values in the more productive Shingobee Lake were −15 per mil and hypolimnion values 

were −19 per mil; epilimnion values in Little Shingobee Lake were −17 per mil and hypolimnion values 

were −20 per mil (R.G. Striegl, USGS, written commun., 2008). Resulting values of δ13C in DIC in the 

epilimnion and hypolimnion of these three lakes were −3 per mil and −5 per mil for Williams,

−8 per mil and −12 per mil for Shingobee, and −10 and −13 for Little Shingobee. Clearly the 

productivity model was working in these three lakes, resulting in the transfer of 13C-depleted organic 

(CO2) and inorganic (DIC) carbon from the epilimnion to the hypolimnion, leaving 13C-enriched carbon 

in the organic and inorganic reservoirs in the epilimnion. The increase in values of δ13C in marl in both 

Williams and Shingobee Lakes in mid-Holocene sediments (fig. 14; Schwalb and Dean, 2002) suggests 

that at least some of this increase in values of δ13C in the DIC reservoirs is due to the productivity 

model. Dean and Stuiver (1993) attributed the sharp increase in values of δ13C in OC in the varved 

sediments of Elk Lake during the prairie period to an increase in production of more hydrogen-rich, 

algal organic matter.

Dean and Schwalb (2002) suggested that some of the mid-Holocene δ13C-enrichment in marl 

and OC in sediments of Williams and Shingobee might be due to the change in the drainage basin from 

the C3-plant-dominated pine forest to C4-plant-dominated prairie grasses having values of δ13C that 

average about −27 and −12 per mil, respectively, a result of their different photosynthetic pathways 

(O’Leary, 1988; Cerling and Quade, 1993; Meyers and Teranes, 2001). The mid-Holocene increase in 

δ13C of OC in Shingobee Lake marl is about 4 per mil, but in Williams Lake marl the increase is about 

10 per mil. Dean and Schwalb (2002) thought that the difference (6 per mil) was due to hydrology; the 

other 4 per mil may represent vegetation change in the drainage basin during the early Holocene. The 

fact that values of δ13C in OC in Shingobee Lake decreased to early Holocene levels following 

reforestation during the late Holocene suggests that vegetation change in the drainage basin and not 

productivity was the main cause of the increase in δ13C in OC in Shingobee Lake in the mid-Holocene, 

and probably some of the increase of 10 per mil in Williams Lake. 

Values of δ13C in marl also increased increase in mid-Holocene sediments in Williams and 

Shingobee Lakes indicating that the entire surface-water carbon reservoir was enriched in 13C. Values of 

δ13C in the shells of Pisidium and ostracodes in Little Shingobee Lake increase in the mid-Holocene 

section from about −9 per mil at 660 cm depth to −7 per mil at 600 cm (about 9 cal ka) and remain high 

at −7 per mil to −8 per mil until 440 cm (about 6.5 c al. ka) (fig. 8). However, unlike in the sediments of 

Williams, Shingobee, and Elk Lakes, values of δ13C in OC in Little Shingobee sediments did not 
34

increase in the mid-Holocene (fig. 9). Values remain between −28 per mil and −30 per mil throughout 

the entire Holocene section. Values of percent OC in Little Shingobee Lake sediments increased from 

about 1 percent to 4 percent during the mid-Holocene, suggesting that productivity had increased, but a 

lack of an increase in δ13C in OC suggests that productivity did not increase. 

The sharp increases in δ13C of marl (fig. 14) and OC (Dean and Schwalb, 2002) in midHolocene sediments in Williams Lake, relative to the more modest increases in marl and OC in 

Shingobee Lake sediments, indicate that the hydrologic closing of Williams Lake had a major role in the 

carbon-isotope signatures in the sediments. As mentioned earlier, values of both δ18O and δ13C in the 

Shingobee River transect of Kendall and others (1997) in the dry late summer increased from 

−11 per mil to −7 per mil, indicating an evaporation trend. Kendall and others concluded that the 

increase in δ13C along the transect reflects hydrologic processes (residence time) and not biological 

recycling (productivity). Therefore, the mid-Holocene increase in δ13C in carbonate shells but not in OC 

in Little Shingobee Lake sediments is likely due to longer residence time during the dry prairie period 

and not due to productivity. Some of the larger mid-Holocene increases in values of δ13C in sediments 

of Shingobee and Elk Lakes are probably also due to hydrology although some of the increases are due 

to productivity, as suggested by the mid-Holocene increase in δ13C in OC in both lakes. Some of the 

mid-Holocene increases in values of δ13C are undoubtedly due to changes from forest to prairie. At this 

time we do not have the data to measure the influence of these three factors (productivity, vegetation 

change, and hydrology) on the isotopic composition of DIC and OC in these lakes, although hydrology 

was obviously a dominant factor in Williams Lake. All three could explain the mid-Holocene increases 

in δ13C in Williams and Shingobee Lakes. 

Pollen Zones

We compare the timing of pollen zones in Little Shingobee Lake with those in Shingobee Lake 

(Locke, 1995), and Elk Lake (Whitlock and others, 1993) in figure 6. The main tie point in the 

chronology of Shingobee Lake is a radiocarbon date of a piece of wood of about 7,700 years ago at the 

horizon of the forest/prairie border. The other pollen zones are extrapolated from dated zone boundaries 

in other lakes in northwestern Minnesota, including Elk Lake. We have calibrated those radiocarbon 

ages reported by Locke by using the online radiocarbon calibration program CalPal. The varve 

timescale in Elk Lake has been well correlated with the radiocarbon timescale (Anderson and others, 

1993; Sprowl, 1993; review by Dean and others, 2002a). For the pollen-zone boundaries in Little 

Shingobee we tried to use the same criteria used by Whitlock and others (1993) for Elk Lake on our 

limited sampling. We would like to see a better correspondence for the prairie period but it is not 

evident in the pollen diagram (fig. 10). The beginning of the prairie period is defined by a decrease in 

Pinus and increases in Quercus, Artemisia, and Gramineae (fig. 10). The end of the prairie period is 

defined by a decrease in Artemisia and an increase in Betula. We are encouraged that the pollen zones in 

Little Shingobee Lake match changes in lithology. Most notable of these changes is the marked increase 

in CaCO3 between 700 cm and 400 cm depth (10 cal. ka and 6 cal. ka (fig. 6). In the white pine zone 

there are marked increases in endogenic CaCO3, OM, Fe(OH)3, and biopal (fig. 6) that correspond with 

the warm Prairie Period as defined in core LSL–B. This increase is probably in response to a slight 

increase in temperature at the beginning of the white pine zone (see discussion below). 

The reconstructed climate of northwestern Minnesota in the Spruce zone was cold and relatively 

dry (Whitlock and others, 1993). Numerical modeling of the Elk Lake pollen data by Bartlein and 

Whitlock (1993) suggests that the temperature in northwestern Minnesota was about 5°C lower in 

January than at present and 2.5° lower in July. Precipitation was about 200 mm less annually than today. 
35

The Jack/Red Pine zone of early Holocene age was slightly drier and warmer than the Spruce 

zone, with closest modern analogs in northern Wisconsin (Whitlock and others, 1993). Reconstruction 

of annual precipitation by comparing the Elk Lake data with modern pollen data shows that the earlymiddle Holocene transition (8.5 to 7.8 cal. ka) is marked by an abrupt decrease in precipitation of about 

200 millimeters per year (mm yr–1) (Bartlein and Whitlock, 1993). Artemisia is the most distinctive 

vegetation of prairie environments west of the Shingobee River headwaters area. An expansion of 

sagebrush pollen in abundance at about 10 cal. ka in Little Shingobee Lake (about 8.7 varve years in Elk 

Lake) marks the initial expansion of the prairie eastward. Reviewing all of the evidence from Elk Lake, 

including pollen, diatoms, ostracodes, geochemistry, mineralogy, and varve thickness, Bradbury and 

others (1993) and Dean and others (2002a) concluded that between about 8.6 and 8.0 cal. ka (on the 

basis of varve counts), Elk Lake went from a well-stratified lake in a pine forest that protected the lake 

from wind mixing, to a well-mixed lake in dry, open prairie and oak savanna receiving wind-blown dust 

as the polar front moved north into northern Canada, and that the most abrupt change occurred at 8.2 

cal. ka within about 100 years. The only other varve-calibrated lake record from Minnesota is from Lake 

of the Clouds in the Boundary Waters Canoe Area of northeastern Minnesota (Craig, 1972). In that 

record Chenopod pollen blown from western Minnesota became important at 8.2 cal. ka. 

The Prairie zone pollen taxa also had abundant Quercus and Gamineae. During this midHolocene period, the forest-prairie border in northwestern Minnesota migrated at least 100 km east of its 

present position, changing from pine forest to sagebrush-and-grass prairie and oak savanna (Whitlock 

and others, 1993). The average annual precipitation during the prairie period was about 100 mm lower 

than at present, and July temperature continued an increasing trend begun in the pine zone to about 2°C 

warmer at the end of the prairie period (Bartlein and Whitlock, 1993). Modern pollen analogs for this 

period are located in the prairies along the North Dakota–Canadian border. Ostracode and diatom data 

from Elk Lake (Forester, 1987; Bradbury and Dieterich-Rurup, 1993) suggest that the early part of the 

mid-Holocene in northwestern Minnesota was cooler throughout the year than at present. 

On the basis of the Elk Lake varve-calibrated record, after 8.2 cal. ka, westerly winds became 

increasingly dominant (Bradbury and others, 1993; Dean and others, 2002a). High concentrations of Al, 

quartz, and plagioclase feldspar indicate an increase in influx of relatively unweathered detrital clastic 

material moved by the wind and forming thicker varves (Dean, 1993). The effects of increased 

temperature and evaporation between 7.2 varve ka and 6.5 varve ka produced the highest salinity in the 

history of Elk Lake, as evidenced by ostracode assemblages (Forester and others, 1987) and presence of 

aragonite needles in the sediments (Dean, 1993). The increase in temperature and evaporation is 

manifested in Little Shingobee Lake sediments by the marked increase in endogenic CaCO3 from about 

20 percent in the pine zone (>700 cm) to 60 percent to 70 percent throughout the prairie period (fig. 5). 

This increase in CaCO3 is the most distinctive feature of the prairie period in Little Shingobee Lake. 

Maxima of eolian indicators in Elk Lake were reached at about 5.8 varve ka. Evidence for increased 

aridity and eolian activity during the mid-Holocene in the Great Plains and Rocky Mountain region is 

abundant, with predominantly westerly or northwesterly winds (see review by Dean and others, 1996). 

Increased productivity in Elk Lake during the prairie period is evidenced by high concentrations 

of algal pigments (Sanger and Hay, 1993), increases in δ13C in both endogenic CaCO3 and hydrogenrich OM (discussed above). Increased productivity in Little Shingobee Lake during the prairie period is 

evidenced by increases in δ13C in CaCO3 of shells (fig. 8) and increases in OM and biopal (fig. 6).

The return of the forest is marked by an increase in arboreal hardwood pollen taxa dominated by 

Quercus and Betula but with significant contributions from Ostrya and Pinus. Precipitation and 

temperature reconstructions show that going into the hardwood zone annual precipitation increased to 

values greater than present, and July temperature remained slightly greater than at present (Bartlein and 
36

Whitlock, 1993). This Hardwood zone has modern analogs in the hardwood-conifer forests of the 

southern Great Lakes region where the climate is wetter than that of the prairie period.

The White Pine zone marks the return of pine forest to northwestern Minnesota, this time 

dominated by white pine (P. strobus) but with important contributions from P. banksiana and P. 

resinosa. White pine reached northeastern Minnesota from the Great Lakes region by about 7.2 cal. ka 

but was unable to spread west during the prairie period (Whitlock and others, 1993). Its appearance 

marks a western retreat of the forest-prairie border characterized by warm, moist conditions in 

northwestern Minnesota. At the beginning of the white pine zone, July temperature was about 1.5°C 

higher than at present, January temperature was about 2°C higher, and precipitation was about 100 mm 

greater (Bartlein and Whitlock, 1993). All three climate variables decreased through the zone to modern 

values. The increase in temperature appears to have been enough to trigger increases in endogenic 

inorganic components in the middle of the White Pine zone in Little Shingobee Lake including CaCO3

and Fe(OH)3 and an increase in productivity as evidenced by increases in endogenic OM and biopal 

(fig. 6). This increase in productivity may also be manifested by an increase in values of δ13C in 

Pisidium (fig. 8), although there are few values. 

Most paleolimnologic parameters in the Elk Lake record remain relatively constant during the 

last 4 ky, which includes the Hardwood and White Pine zones. Bradbury and others (1993) concluded 

that overall the last 4 ky of the Elk Lake record represents the establishment of modern climatic and 

environmental regimes. However, there are fluctuations of some paleolimnologic parameters that may 

be associated with decadal to centennial variability in modern climates, such as the Little Ice Age 

(Anderson, 1993; Dean, 2002). Fluctuations in many paleolimnologic parameters in the last 1,500 years 

of Elk Lake history are documented by Bradbury and others (2002), Dean (2002), and Dean and others, 

(2002b). 

Conclusions

1. The major components in the Holocene sediments in both Little Shingobee Lake and Little 

Shingobee Fen are allogenic detrital clastic material (aluminosilicates), endogenic CaCO3, and 

organic matter. The main difference is that in Little Shingobee Fen CaCO3 is by far the dominant 

component with little change through the Holocene section with an average of 61 percent

CaCO3. Shingobee Lake sediments are dominated by detrital material (7 percent to 82 percent) 

interrupted by several major pulses of CaCO3 that diluted the detrital material. The lack of 

detrital material in Little Shingobee Fen (average of 4.4 percent) may have been because the sill 

that separated Little Shingobee Lake and the lake that occupied the basin under the present-day 

fen restricted inflow of detrital material. The organic matter content in Holocene sediments of 

Little Shingobee Fen (average 13 percent) is about twice that of Little Shingobee Lake sediments 

(6 percent). 

2. An early Holocene carbonate pulse resulted in the accumulation of about 80 percent CaCO3 in 

the sediments of both Little Shingobee Lake and Fen and also in the sediments of nearby 

Williams and Shingobee Lakes. This pulse of CaCO3 deposition was likely the result of 

postglacial leaching of CaCO3 from the calcareous glacial drift in the drainage basin into the 

lakes that triggered endogenic calcite precipitation.

3. Almost all of the Fe and Mn in the sediments in both Little Shingobee Lake and Fen is 

endogenic, even in the late glacial detrital sediments, and is assumed to reside in oxyhydroxide 

minerals. The source of high concentrations of Fe and Mn that accumulate in the hypolimnion of 

Shingobee Lake today is not shallow groundwater but rather is a deep aquifer that overlies 
37

Cretaceous bedrock. Therefore, it seems likely that the source of Fe and Mn that precipitated as 

endogenic minerals that accumulated in the sediments of Little Shingobee Lake and Little 

Shingobee Fen, as well as those in Williams and Shingobee Lakes, was postglacial activation of 

that deep aquifer.

4. Values of δ18O in carbonate shells in Little Shingobee Lake sediments are remarkably constant 

throughout the Holocene, indicating that environmental conditions affecting oxygen-isotope 

composition were fairly constant. However, there is an increase in δ13C in mid-Holocene in 

carbonate shells but not in organic carbon, and this increase is likely due to longer residence 

time during the dry prairie period. 

5. The most striking change in the vegetation of Shingobee River headwater area, and northwestern 

Minnesota in general, is the transition from late Pleistocene and early Holocene forests of spruce 

and then pine to a mid-Holocene open sagebrush-oak savanna. Forests returned to northern 

Minnesota during the late Holocene, first as a hardwood forest followed by a return to a pine 

forest.

Acknowledgments

This work was funded by the U.S. Geological Survey Global Change and Climate History and 

Earth Surface Dynamics Programs. We thank Lesleigh Anderson, Don Rosenberry, and Tom Winter for 

helpful comments on an earlier version of this paper. Cores are archived at LacCore (National 

Lacustrine Core Repository), Department of Geology and Geophysics, University of Minnesota—Twin 

Cities.

References Cited

Anderson, R.Y., 1993, The varve chronometer in Elk Lake—Record of climatic variability and evidence 

for solar/geomagnetic14C-climate connection, in Bradbury, J.P., and Dean, W.E., eds., Elk Lake, 

Minnesota—Evidence for rapid climate change in the north-central United States: Geological Society 

of America Special Paper 276, p. 46–68. 

Anderson, R.Y, Bradbury, J.P., Dean, W.E., and Stuiver, Minze, 1993, The varve chronology of Elk 

Lake sediments—Coring, sampling, and time-series construction, in Bradbury, J.P., and Dean, W.E., 

eds., Elk Lake, Minnesota—Evidence for rapid climate change in the north-central United States: 

Geological Society of America Special Paper 276, p. 37–44. 

Bartlein, P.J., and Whitlock, Cathy, 1993, Paleoclimatic interpretation of the Elk Lake pollen record, in

Bradbury, J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence for rapid climate change in the 

north-central United States: Geological Society of America Special Paper 276, p. 275–294. 

Bernasconi, S.M., Barbieri, Alberto, and Simona, Marco, 1997, Carbon and oxygen isotope variations in 

sedimenting organic matter in Lake Lugano: Limnology and Oceanography, v. 42, p. 1755–1765. 

Bradbury, J.P., Cumming, B.E., and Laird, K.A., 2002, A 1,500-year record of climatic and 

environmental change in Elk Lake, Minnesota—III. Measures of past primary productivity: Journal of 

Paleolimnology, v. 27, p. 320–331. 

Bradbury, J.P., Dean, W.E., and Anderson,R.Y., 1993, Holocene climatic and limnologic history of the 

north-central United States as recorded in the varve sediments of Elk Lake, Minnesota—A synthesis, 

in Bradbury, J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence for rapid climate change in 

the north-central United States: Geological Society of America Special Paper 276, p. 309–328. 
38

Bradbury, J.P., and Dieterich-Rurup, Katherine, 1993, Holocene diatom paleolimnology of Elk Lake, 

Minnesota, in Bradbury, J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence for rapid climate 

change in the north-central United States: Geological Society of America Special Paper 276, p. 215–

238. 

Carter, Virginia, Gammon, P.T., and Rybicki, N.B., 1997, Vegetation of Little Shingobee Fen, in

Winter, T.C., ed., Interdisciplinary Research Initiative—Hydrological and biogeochemical research in 

the Shingobee River headwaters area, north-central Minnesota: U. S. Geological Survey Water 

Resources Investigations Research 96–4215, p. 161–166. 

Cerling, T.E., and Quade, Jay, 1993, Stable carbon and oxygen isotopes in soil carbonates, in Swart, 

P.K., Lohmann, K.C., McKenzie, Judith, and Savin, Samuel, eds., Climate change in continental 

isotopic records: American Geophysical Union Geophysical Monograph 78, p. 217–231. 

Craig, A.J., 1972, Pollen influx to laminated sediments—A pollen diagram from northeastern 

Minnesota: Ecology, v. 53, p. 46–57. 

Davison, W., and Seed, G., 1983, The kinetics of oxidation of ferrous iron in synthetic and natural 

waters: Geochimica et Cosmochimica Acta, v. 47, p. 67–79. 

Dean, W.E., Jr., 1974, Determination of carbonate and organic matter in calcareous sediments and 

sedimentary rocks by loss on ignition—Comparison with other methods: Journal of Sedimentary 

Petrology, v. 44, p. 242–248. 

Dean, W.E., 1993, Physical properties, mineralogy, and geochemistry of Holocene varved sediments 

from Elk Lake, Minnesota, in Bradbury, J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence 

for rapid climate change in the north-central United States: Geological Society of America Special 

Paper 276, p. 135–158. 

Dean, W.E., 2002, A 1,500-year record of climatic and environmental change in Elk Lake, Minnesota—

II. Geochemistry, mineralogy, and stable isotopes: Journal of Paleolimnology, v. 27, p. 301–319. 

Dean, W.E., Ahlbrandt, T.S., Anderson, R.Y., and Bradbury, J.P., 1996, Regional aridity in North 

America during the middle Holocene: The Holocene, v. 6, p. 145–155. 

Dean, W.E., Anderson, R.Y., Bradbury, J.P., and Anderson, D., 2002b, A 1,500-year record of climatic 

and environmental change in Elk Lake, Minnesota I—Varve thickness and gray-scale density: Journal 

of Paleolimnology, v. 27, p. 287–299. 

Dean, W.E., Arthur, M.A., and Claypool, G.E., 1986, Depletion of 13C in Cretaceous marine organic 

matter—Source, diagenetic, or environmental signal?: Marine Geology, v. 70, p. 119–158. 

Dean, W.E., and Bradbury, J.P., 1997, Transects of organic carbon, calcium carbonate, and diatoms in 

surface sediments of Williams and Shingobee Lake, in Winter, T.C., ed., Interdisciplinary Research 

Initiative—Hydrological and biogeochemical research in the Shingobee River headwaters area, northcentral Minnesota: U.S. Geological Survey Water Resources Investigations Research 96–4215, p. 

117–130. 

Dean, W.E., Forester, R.M., and Bradbury, J.P., 2002a, Early Holocene change in atmospheric 

circulation in the Northern Great Plains—An upstream view of the 8.2 ka cold event: Quaternary 

Science Reviews, v. 21, p. 1763–1775. 

Dean, W.E., Neff, B., Rosenberry, D., Winter, T.C., and Parkhurst, R., 2003, The significance of ground 

water to the accumulation of iron and manganese in the sediments of two hydrologically distinct lakes 

in north-central Minnesota—A geological perspective: Ground Water, v. 41, p. 951–963. 

Dean, W.E., and Schwalb, Antje, 2002, The lacustrine carbon cycle as illuminated by the waters and 

sediments of two hydrologically distinct hardwater lakes in northwestern Minnesota: Journal of 

Sedimentary Research, v. 72, p. 416–431. 
39

Dean, W.E., and Stuiver, Minze, 1993, Stable carbon and oxygen isotope studies of the sediments of Elk 

Lake, Minnesota, in Bradbury, J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence for rapid 

climate change in the north-central United States: Geological Society of America Special Paper 276, 

p. 163–180. 

Engleman, E.E., Jackson, L.L., Norton, D.R., and Fischer, A.G., 1985, Determination of carbonate 

carbon in geological materials by coulometric titration: Chemical Geology, v. 53, p. 125–128. 

Faegri, Knut, Kaland, P.E., Krzywinski, Knut, 1989, Textbook of pollen analysis, 4th ed.: New York, 

Blackwell, 328 p. 

Forester, R.M., 1987, Late Quaternary paleoclimate records from lacustrine ostracodes, in Ruddiman, 

W.F., and Wright, H.E., Jr., eds., The geology of North America, v. K–3, chapter 12, North America 

and adjacent oceans during the last deglaciation: Geological Society of America, p. 261–276. 

Gaugh, H.G., Jr., 1982, Multivariate analysis in community ecology: New York, Cambridge University 

Press, 298 p. 

Gorham, Eville, and Sanger, J.E., 1975, Fossil pigments in Minnesota lake sediments and their bearing 

upon the balance between terrestrial and inputs to sedimentary organic matter: Verhandlungen 

International Vereinigun für Limnologie, v. 19, p. 267–273. 

Grimm, Eric, 1987, CONISS—A Fortran 77 program for stratigraphically constrained cluster analysis 

by the method of incremental sum of squares: Computers and Geosciences, v. 13, p. 13–35. 

Hollander, D.J., and McKenzie, J.A., 1991, CO2 control on carbon-isotope fractionation during aqueous 

photosynthesis—A paleo-pCO2 barometer: Geology, v. 19, p. 929–932. 

Hollander, D.J., and Smith, M.A., 2001, Microbially mediated carbon cycling as a control on the δ13C of 

sedimentary carbon in eutrophic Lake Mendota (USA)—New models for interpreting isotopic 

excursions in the sedimentary record: Geochimica et Cosmochimica Acta, v. 65, p. 4321–4337. 

Kendall, Carol, Evans, W.C., Reddy, M.M., and Schuster, P.F., 1997, Application of stable isotopes in 

the Shingobee River headwaters area, in Winter, T.C., ed., Interdisciplinary Research Initiative—

Hydrological and biogeochemical research in the Shingobee River Headwaters Area, north-central 

Minnesota: U.S. Geological Survey Water Resources Investigations Research 96–4215, p. 63–70. 

LaBaugh, J.W., 1997, The plankton communities of Shingobee and Williams Lakes, in Winter, T.C., 

ed., Interdisciplinary Research Initiative—Hydrological and biogeochemical research in the 

Shingobee River Headwaters Area, north-central Minnesota: U.S. Geological Survey Water 

Resources Investigations Research 96–4215, p. 97–104. 

Locke, S.M., 1995, A paleohydrologic model applied to the Holocene sediment stratigraphy of two 

lakes in north-central Minnesota: Minneapolis, University of Minnesota, Ph.D. dissertation, 258 p. 

Locke, S.M., and Schwalb, Antje, 1997, Sediment stratigraphy and paleolimnological characteristics of 

Williams and Shingobee Lakes, in Winter, T.C., ed., Interdisciplinary Research Initiative—

Hydrological and biogeochemical research in the Shingobee River Headwaters Area, north-central 

Minnesota: U.S. Geological Survey Water Resources Investigations Research 96–4215, p. 187–192. 

McKenzie, J.A., and Hollander, D.J., 1993, Oxygen-isotope record in recent carbonate sediments from 

Lake Greifen, Switzerland (1750–1986)—Application of continental isotopic indicator for evaluation 

of changes in climate and atmospheric circulations patterns, in Swart, P.K., Lohman, K.C., McKenzie, 

J.A., and Savin, Samuel, eds., Climate change in continental isotope records: American Geophysical 

Union Monograph 78, p. 101–111. 

Megard, R.O., Bradbury, J.P., and Dean, W.E., 1993, Climatic and limnologic setting of Elk Lake, in

Bradbury, J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence for rapid climate change in the 

north-central United States: Geological Society of America Special Paper 276, p. 19–36. 
40

Meyers, P.A., and Teranes, J.L., 2001, Sediment organic matter, in Last, W.M., and Smol, J.P., eds., 

Tracking environmental change using lake sediments, v. 2: Dordrecht, The Netherlands, Kluwer 

Academic Publishers, p. 239–269. 

Moore, D.M., and Reynolds, R.C., Jr., 1989, X-ray diffraction and identification and analysis of clay 

minerals: Oxford, U.K., Oxford University Press, 332 p. 

Moores, H.D., and Norton, A.R., 1997, Glacial landscape evolution of the Itasca/St. Croix moraine 

interlobate area including the Shingobee River Headwaters Area, in Winter, T.C., ed., 

Interdisciplinary Research Initiative—Hydrological and biogeochemical research in the Shingobee 

River Headwaters Area, north-central Minnesota: U.S. Geological Survey Water Resources 

Investigations Research 96–4215, p. 11–18. 

Myrbo, Amy, and Shapley, M.D., 2006, Seasonal water-column dynamics of dissolved inorganic carbon 

stable isotope compositions (δ13CDIC) in small hardwater lakes in Minnesota and Montana: 

Geochimica et Cosmochimica Acta, v. 70, p. 2699–2714. 

O’Leary, M.H., 1988, Carbon isotopes in photosynthesis: BioScience, v. 38, p. 328–336. 

Overpeck, J.T., Webb, Thompson, III, and Prentice, I.C., 1985, Quantitative interpretation of fossil 

pollen spectra—Dissimilarity coefficients and the method of modern analogs: Quaternary Research, v. 

23, p. 87–108. 

Rosenberry, D.O., Winter, T.C., Merk, D.A., Leavesley, G.H., and Beaver, L.D., 1997, Hydrology of 

the Shingobee River Headwaters Area, in Winter, T.C., ed., Interdisciplinary Research Initiative—

Hydrological and biogeochemical research in the Shingobee River headwaters area, north-central 

Minnesota: U. S. Geological Survey Water Resources Investigations Research 96–4215, p. 19–24.

Sanger, J.E., and Hay, R.J., 1993, Fossil pigments in Holocene varved sediments, Elk Lake, Minnesota, 

in Bradbury, J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence for rapid climate change in 

the north-central United States: Geological Society of America Special Paper 276, p. 181–188. 

Schwalb, Antje, and Dean, W.E., 2002, Reconstruction of hydrological changes and effective moisture 

from north-central USA lake sediments: Quaternary Science Reviews, v. 21, p. 1541–1554. 

Schwalb, Antje, Locke, S.M., and Dean, W.E., 1995, Ostracode δ18O and δ13C evidence of Holocene 

environmental changes in the sediments of two Minnesota lakes: Journal of Paleolimnology, v. 14, p. 

281–296. 

Stauffer, R.E., and Armstrong, D.E., 1986, Cycling of iron, manganese, silica, phosphorus, calcium and 

potassium in two stratified basins of Shagawa Lake, Minnesota: Geochimica et Cosmochimica Acta, 

v. 50, p. 215–229. 

Striegl, R.G., Kortelainen, Pirkko, Chanton, J.P., Wickland, K.P., Bugna, G.C., and Rantakari, Miitta, 

2001, Carbon dioxide partial pressure and 13C content of north temperate and boreal lakes at spring 

ice melt: Limnology and Oceanography, v. 46, p. 941–945. 

Striegl, R.G., and Michmerhuizen, C.M., 1998, Hydrologic influence on methane and carbon dioxide 

dynamics in two north-central Minnesota lakes: Limnology and Oceanography, v. 43, p. 1519–1529. 

Stuiver, Minze, Reimer, P.J., Bard, Edouard, Beck, J.W., Burr, G.S., Hughen, K.A., Kromer, Bernd, 

McCormac, Gerry, Van der Plicht, Johannes, and Spurk, Marco, 1998, INTCAL98 radiocarbon age 

calibration, 24,000–0 cal BP: Radiocarbon, v. 40, p. 1041–1083. 

Sprowl, D.R., 1993, On the precision of the Elk Lake varve chronology, in Bradbury, J.P., and Dean, 

W.E., eds., Elk Lake, Minnesota—Evidence for rapid climate change in the north-central United 

States: Geological Society of America Special Paper 276, p. 69–74. 

Tian, Jian, Brown, T.A., and Hu, F.S., 2005, Comparison of varve and 14C chronologies from Steel 

Lake, Minnesota, USA: The Holocene, v. 15, p. 510–517. 
41

Tipping, E., Woof, C., and Cooke, D., 1981, Iron oxide from a seasonally anoxic lake: Geochimica et 

Cosmochimica Acta, v. 45, p. 1441–1419. 

Von Grafenstein, Ulrich, Erlenkeuser, Helmut, and Trimborn, Peter, 1999, Oxygen and carbon isotopes 

in modern fresh-water ostracod valves—Assessing vital offsets and autoecological effects of interest 

for palaeoclimate studies: Palaeogeography, Palaeoclimatology, and Palaeoecology, v. 148, p. 133–

152. 

Wetzel, R.G., 2001, Limnology—Lake and river ecosystems (3d ed): San Diego, Calif., Academic 

Press, 1006 p. 

Whitlock, Cathy, Bartlein, P.J., and Watts, W.A., 1993, Vegetation history of Elk Lake, in Bradbury, 

J.P., and Dean, W.E., eds., Elk Lake, Minnesota—Evidence for rapid climate change in the northcentral United States: Geological Society of America Special Paper 276, p. 251–274. 

Winter, T.C., ed., 1997, Interdisciplinary Research Initiative—Hydrological and biogeochemical 

research in the Shingobee River Headwaters Area, north-central Minnesota: U. S. Geological Survey 

Water Resources Investigations Research 96–4215, 210 p. 

Wright, H.E., Jr., 1976, The dynamic nature of Holocene vegetation: Quaternary Research, v. 24, p. 

115–120. 

Wright, H.E., Jr., Stefanova, Ivana, Tian, Jian, Brown, T.A., and Hu, F.S., 2004, A chronological 

framework for the Holocene vegetational history of central Minnesota—The Steel Lake pollen record: 

Quaternary Science Reviews, v. 23, p. 611–626. 

Xia, J., Ito, Emi, and Engstrom, D.R., 1997a, Geochemistry of ostracode calcite—An experimental 

determination of oxygen isotope fractionation: Geochimica et Cosmochimica Acta, v. 61, p. 377–382. 

Xia, J., Ito, Emi, and Engstrom, D.R., 1997b, Variation of stable-isotope and trace-element composition 

of lacustrine ostracodes from the northern Great Plains, part 2—The effects of lake water chemistry 

and seasonal temperature variation on Candona rawsoni: Geochimica et Cosmochimica Acta, v. 61, p.

383–391.