Gorin Thesis Research

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Paleoclimate Reconstruction from a Weybridge Cave Speleothem, Vermont
By
Drew Gorin
Submitted in partial fulfillment of
The requirements for the degree of
Bachelor of Arts
Department of Geology
Middlebury College
Middlebury, Vermont
May 2016

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Gorin, Andrew L., 2016. Paleoclimate Reconstruction from a Weybridge Cave
Speleothem, Vermont: Unpublished Senior Thesis, Middlebury College, Middlebury,
VT, 72p.
Abstract
Understanding future climate change may prove to be one of the most important
scientific endeavors of this century. Studying past climate change allows scientists to
design global models that better simulate and predict future climate patterns. This
project focuses on the climate of Weybridge and the surrounding area over the past 5,000
years by studying the geochemistry of a speleothem taken from Weybridge Cave. This
involves two primary tasks. The speleothem itself was sampled and analyzed for stable
isotopes of oxygen and carbon, and a hydrology study was conducted on the cave. The
speleothem oxygen isotope chemistry provides insight into past precipitation quantity,
source, and possibly temperature, while the carbon isotope chemistry provides insight
into vegetation changes. The speleothem was dated with U-series dating, and spans
between 4.8-1.5ka BP.
My data suggests that δ18
O fluctuations have historically been controlled by
precipitation source and changes in seasonal precipitation distribution. As the precession
orbital variation has created warmer northern hemisphere winters, changes in the North
American Winter-Vortex and in the Bermuda High have profoundly influenced the
precipitation balance in the region. My data also record the Middle-Holocene
Transition, or 4.2 ka event, which was a time of profound global climate change. This
study is of particular interest because there have been very few speleothem paleoclimate
reconstructions done in the Northeastern United States.

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Acknowledgements
I would like to thank my advisors Jeff Munroe and Will Amidon for their
unwavering support and wisdom throughout this project. The two of them have truly
taken me under their wing over the last four years. Thanks to the both of you for letting
me join in on your field projects, and for taking the time to give me a deeper
understanding of what it means to be a geologist. The countless opportunities afforded to
me by the Middlebury Geology Department have instilled in me a true passion for
geological inquiry and research.
Thanks also to David Gilliken, at Union College, for your technical help, and for
access to your facilities. Jamie Shanley at the USGS in Montpelier also provided
invaluable data for this project.
I’m obviously eternally indebted to my parents and family for their unconditional
love and support throughout my time at Middlebury. Their guidance is what allowed me
to make it to Middlebury in the first place, and their continued support has kept me
grounded throughout my college years.
Finally, I would like to thank the Middlebury Undergraduate Research Office for
their financial support of this project, and of other independent research projects that I’ve
been involved in.

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For J-bro, I’m sorry I stole your major,

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Table of Contents
Introduction 8
Background 9
Geologic Setting............................................................................................................ 11
Speleothem Stable Isotope Geochemistry......................................................................... 14
δ18
O Values from Speleothems.................................................................................... 15
δ13
C values from Speleothems..................................................................................... 18
Hendy Test................................................................................................................ 19
Previous Isotopic Work in The Northeastern United States ................................................ 20
Paleoclimate History of the Northeastern United States..................................................... 23
Speleothem Uranium-Thorium Dating............................................................................. 25
Speleothem Physical Appearance.................................................................................... 28
Concluding Notes .......................................................................................................... 29
Methods 30
Speleothem Sampling Strategy........................................................................................ 30
Scanning Electron Microscopy........................................................................................ 31
Water Sampling............................................................................................................. 31
Results 33
Age Model.................................................................................................................... 33
Stable Isotope Data........................................................................................................ 37
Hendy Tests .................................................................................................................. 40
Modern δ18
O values in Vermont...................................................................................... 42
Weybridge Cave Drip Water........................................................................................... 43
Scanning Electron Microscope Data ................................................................................ 45
Discussion 50
Modern Vermont Climate............................................................................................... 50
The North Atlantic Subtropical Anticyclone..................................................................... 52
The North American Winter-Vortex ................................................................................ 53
δ18
O Interpretation ......................................................................................................... 54
δ13
C Interpretation ......................................................................................................... 60
Conclusion 60
Works Cited 62

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Table of Figures
Figure 1: Weybridge Cave Location 12
Figure 2: Lidar Image of Area Surrounding Weybridge Cave 13
Figure 3: Map of Weybridge Cave 13
Figure 4: Local Meteoric Water Line from Underhill, VT 17
Figure 5: 238
U decay Series 26
Figure 6: Diagram of Applied Detrital Th Corrections 27
Figure 7: Weybridge Cave Speleothem photo with U-Th ages 28
Figure 8: Growth Hiatus and Isotope Transect Diagram 30
Figure 9: Photos Showing Water Sampling Methods 32
Figure 10: Initial Speleothem Age Model 34
Figure 11: Refined Speleothem Age Model 36
Figure 12: δ18
O and δ13
C Stable Isotope Data 37
Figures 13 and 14: Raw δ18
O Values With Zoomed Windows 39
Figure 15: δ18
O and δ13
C Speleothem Correlation Diagram 40
Figure 16: Hendy Test Non-Correlation Diagrams 41
Figure 17: The Modern Seasonality of δ18
O in St. Johnsbury, VT 43
Figure 18: Weybridge Cave Drip Water and Precipitation δ18
O 44
Figure 19: SEM photos of Speleothem Detrital Layers 47
Figure 20: EDS Images Showing Elemental Composition of Detrital Layers 48
Figure 21: Box Plots of the Distribution of δ18
O Values for Detrital and Non-Detrital
Layers in the Speleothem 49
Figure 22: Schematic Diagram of Modern VT Precipitation Sources 52
Figure 23: Potential Track of Retracting North American Winter-Vortex 53
Figure 24: Holocene Insolation Change 57
Figure 25: δ18
O and Solar Insolation Correlation 58

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List of Tables
Table 1: Correlations between δ18
O and δ13
C in different speleothem growth regions 38
Table 2: Average St. Johnsbury Precipitation δ18
O 51
Appendix 1: Raw δ18
O and δ13
C Data 70

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Introduction
Understanding future climate change may prove to be one of the most important
scientific endeavors of this century. Due to the long timespans over which climate
typically varies, one needs to study past climate change in order to make accurate
predictions about the future. This project seeks to contribute to the base of knowledge
about paleoclimate change in Vermont by conducting a study on a speleothem collected
from Weybridge Cave.
The term speleothem refers to calcareous cave growths that form over time from
cave drip water. The concept that speleothems can be analyzed to produce paleoclimate
data originally stems from Hendy’s 1968 study, which quantified the relationship
between precipitation quantity, temperature, and oxygen isotope ratios in cave carbonates
(Hendy, 1968). Studying localized paleoclimate change at high resolutions was
historically possible only in polar regions where ice cores could be obtained from
glaciers. These records have been useful, but they often only provide data regarding how
climate has changed at extreme latitudes. More recently, techniques have been developed
to produce high-resolution climate records from temperate regions. Cave speleothems,
such as stalactites or stalagmites are a primary example of this. This presents a unique
opportunity to the scientific community because these features are found at all latitudes,
which presents opportunities for close study of more temperate regions. For the last
decade or so, scientists around the world have reconstructed local climate using this
technique, however, very few similar studies have been conducted in the Northeastern
United States (McDermott, 2003., Fairchild, 2006., and Lachniet, 2009).

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This study aims to produce and interpret a high-resolution Vermont
paleoclimate reconstruction using a speleothem collected in 2013 from
Weybridge Cave. This study involves two primary tasks. First, geochemical data
from the speleothem itself must be processed. This involves physically drilling
samples from the speleothem and employing geochemical instrumental analysis
techniques, which will be discussed later. Second, the modern hydrology of the
cave and surrounding region must be understood. This is due to the nature of
speleogenic processes, which involve both precipitation and groundwater
transportation.
Background
A thorough understanding of karst landscape dynamics is key to studying
speleothems. Most often, these landscapes are made of limestone and dolostone bedrock.
Karst landscape evolution is dominated by the chemical dissolution of limestone —
CaCO3, which results in the creation of many distinct features such as ridges, towers,
fissures, sinkholes and caves (White, 2007). These features are constantly being created
and modified by the interaction of carbonic acid-enriched groundwater with the highly
soluble bedrock. A simplified version of this reaction can be represented by equation 1
(White, 2007).
(1) H2O + CaCO3(s) + CO2(aq) <—> Ca2+
(aq) + HCO3
-
(aq)

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Understanding this reaction is critical to understanding how speleothems form,
and why they are such precise recorders of paleoclimate. The water in the reaction is
sourced from local precipitation, the calcium carbonate is dissolved from limestone
bedrock, and the carbon is primarily sourced from local plant respiration and
decomposing organic matter (White, 2007). While one might intuitively assume that a
portion of CO2 must come from atmospheric contributions, most rainwater only contains
~0.0037% carbonate. In contrast, by the time rainwater percolates through organic rich
soils, it may contain as much as 10% CO3
2-
(White, 2007). The groundwater then moves
through the limestone through joints, cracks and small pours until it enters a cave through
the ceiling. Because the soil atmosphere has such a high concentration of CO2, carbonate
is dissolved from the bedrock in the karst. This process is facilitated by the fact that high
amounts of carbon dioxide in the soil atmosphere lead to a lowering of soil water pH due
to the production of carbonic acid. Once the drip water enters the cave, CO2 is lost to the
cave environment via degassing, which raises the pH and forces the precipitation of
carbonate.
Due to the nature of this process, more CaCO3 is deposited in the summer months.
This is because far more respiration occurs in the summer than in the winter. Respiration
is directly and indirectly responsible for producing CO2, carbonic acid, and bicarbonate in
the soil horizon. These chemicals drive the disparity in pCO2 between the cave and
groundwater environments, which is what controls speleothem deposition (White, 2007).
Without this pCO2 disparity, not enough CaCO3 is dissolved in solution from the bedrock
to cause speleothem deposition in the cave.

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By definition, speleothems grow through the above process, which is also
responsible for the laminations commonly seen in speleothems. It is because of these
consistent processes, and because of the simple nature of speleothem chemistry that they
are such valuable paleoclimate archives.
Geologic Setting
Weybridge Cave is located within the Beldens Member of the Chipman formation
(Figures 1, 2, and 3) which is a fine grained, Ordovician limestone (Ratcliffe et al.,
2011). The rocks show light-gray to creamy-white weathering and sometimes contain
orange-weathering dolostone and reddish or hematitic calcite. The section of the Beldens
Member directly surrounding the cave is composed primarily of fine-grained limestone
(Ratcliffe, 2011). Weybridge Cave is Vermont’s second longest solutional cave and has
a surveyed length of 458 m and almost 40 m of relief (Quick, 2012). The beds strike at
N85°E and dip at about 14°SE (Long, 1996).
Long (1996) speculates on the formational history of Weybridge Cave. He
suggests that it began as a phreatic cave forming along the strike of the beds. Caves like
this form by inundation of flowing groundwater, which causes dissolution of limestone
over time via equation 1. Weybridge Cave began its life in a phreatic environment
completely under water (Long, 1996). Over time, chemical equilibrium and kinetics are
the limiting factors in the rate of cave formation, which can be as fast as 0.01 cm/y to 0.1
cm/y. This suggests that fractures, joints or pores may take on the order of 104
to 105
years to reach a size large enough for human entry (Palmer, 1991).

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Figure 1: Weybridge Cave Location (Bedrock map data from Suzanne W. Nicholson,
2007)
^
VCGI
438000
438000
440000
440000
442000
442000
444000
444000
446000
446000
448000
448000
450000
450000
167000
167000
170000
170000
173000
173000
176000
176000
179000
179000
182000
182000
Bedrock Type
biotite gneiss
black shale
dolostone (dolomite)
limestone
marble
mica schist
quartzite
slate
Weybridge Cave
0 5
Kilometers

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Figure 2: LiDAR Image of Area Surrounding Weybridge Cave (Perzan, 2014). Arrow
marks cave entrance.
Figure 3: Map of Weybridge Cave (Quick, 2012).
Cave Entrance

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Speleothem Stable Isotope Geochemistry
Introduction
Two predominant techniques are used when trying to recreate paleoclimate using
cave speleothems. Scientists typically study δ18
O and δ13
C values from a transect of drill
holes that runs perpendicular to the laminations in the speleothem, or the growth axis.
These techniques became prevalent after advances in mass spectrometry in the early 90s
allowed for precise U-Th dating on smaller sample sizes (Wong, 2015). Because the
CaCO3 in speleothems is delivered to caves by groundwater, the oxygen isotope
compositions from the water is proportional to precipitation oxygen isotope composition.
Oxygen isotope variability is typically represented in delta notation, which is expressed
as δ18
O shown in equation 2.
(2)
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∗ 1000
The standards are universally agreed upon water samples to which all other samples are
compared. SMOW, a common standard, stands for Standard Mean Ocean Water, and is
most often used for calculating δ18
O values (Fricke, 1999). In speleothems, this ratio is
thought to be proportional to temperature at the time of deposition, which is the key
principle exploited by the technique. Many other factors contribute to δ18
O values, which
will be summarized later. δ13
C, which will also be reviewed later, is the second most
commonly used speleothem paleoclimate proxy. This value is determined by equation 3.

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(3)
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!"#$%#&%
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∗ 1000
The simplified interpretation of this value is that it represents the track of photosynthesis
used by the primary biomass at the time of carbonate deposition. The samples usually
use the Pee Dee Belemnite as a standard.
A less commonly used isotope ratio is δD. This time, the ratio is between two
hydrogen isotopes. The heavier hydrogen isotope is referred to as Deuterium, hence δD.
The equation used to calculate this value is shown in equation 4. Like δ18
O Values,
deuterium isotope ratios are also compared to SMOW.
(4)
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!!"#$%#&%
∗ 1000
δ18
O Values from Speleothems
Generating δ18
O from cave speleothems produces a wealth of paleoclimate
information, which is best interpreted in the context of additional information about the
hydrology and climatic setting of the cave. The most vital assumption made in this
measurement is that the δ18
O value of speleothem carbonate is related to just two
variables: the δ18
O value of the cave drip water and the cave temperature (Lachniet,
2008). Because cave temperatures are usually equal to the mean annual surface
temperature, and because oxygen isotope fractionation is temperature dependent,

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scientists use this value as a paleothermometer. Unfortunately, variations in the global
water cycle fractionate oxygen isotopes in different ways, which often dominate the
speleothem isotope signal, making it difficult to interpret.
Because of the Dole effect, 18
O and 16
O are not uniformly distributed around the
world (Hardt, 2007). The Dole effect states that evaporation disproportionately removes
16
O from bodies of water, which changes their δ18
O values depending on the amount of
evaporation. The rainout effect, which also affects this distribution, states that a
precipitation system delivers water with varying δ18
O values throughout its life. This is
because 18
O is heavier, and therefore more likely to rain out sooner in the system’s
lifespan. In order to see past these such alterations, scientists attempt to identify and
quantify the confounding variables affecting δ18
O. This is accomplished by a detailed
hydrological study of the cave and surrounding region.
First, for each study, a local meteoric water line must be determined (Figure 4).
This includes measuring δ18
O and δD of local precipitation over an extended period of
time.

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This approach allows the investigator to understand how the oxygen and
hydrogen isotope concentrations vary within that particular region. Creating a local
meteoric water line, and comparing it to the global meteoric water line is important for
contextualizing δ18
O data. Generating this data makes the evaluation of soil and drip
water evaporation relative to precipitation possible. It also helps to constrain the seasonal
contributions to drip waters, and helps to estimate moisture recycling (Lachniet, 2008).
A LMWL was created about 45 miles northeast of Weybridge Cave in 2000, which will
be examined in the interpretations made from this study (Abbott et al., 2000).
Figure 4: Local Meteoric Water Line from Underhill, VT
(Abbot et al., 2000)

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Creating a LMWL also helps to contextualize the unequal geographic distribution
of δ18
O ratios. A few effects cause this unequal distribution. The altitude effect involves
a decrease in δ18
O values as altitude increases (Clark and Fritz, 1997). The continental
effect states that δ18
O values decrease with increased distance from the ocean
(Dansgaard, 1964). The last relevant factor is called the amount effect, which states that
there is a decreased δ18
O with increased rainfall quantity within a given storm
(Dansgaard, 1964).
The transit times of water in the vadose zones are also key to interpreting δ18
O
values. This is quantified by measuring the lag between δ18
Oprecipitation values and δ18
Ocave
drip-water values. Two useful end members for visualizing this concept are a cave with a
thin roof that allows water to permeate into the cave quickly, and a cave with a thick roof
where the water has a long residence time before entering the cave. Given a high enough
carbonate accumulation rate, a speleothem from a cave with short groundwater residence
time would best record rapid, high-frequency climate events (McDonald et al., 2007),
while a speleothem from a cave with long groundwater residence time best records long
term average climate signals.
δ13
C values from Speleothems
Although the scientific community is constantly refining their knowledge of δ13
C
as a paleoclimate proxy, there are still many confounding effects on δ13
C values. As
previously stated, the canonical view is that speleothem carbon originates in dissolved
organic carbon found in pore spaces in the soil (Wong, 2015). The carbon is then
dissolved into groundwater and slowly transported into the karst. However, a number of

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other factors that affect δ13
C have recently been discovered. For instance, it is suggested
that there is an inherent bias towards C3 photosynthesis values because tree roots (C3
plants) often reach a depth close to caves. The scientific community has also called into
the question the effect of fluctuating pCO2 (Schubert and Jahren, 2012). It seems that the
partial pressure of carbon dioxide affects the magnitude of carbon isotope fractionation in
C3 plants. The average regional humidity also affects the magnitude by which C3 plants
fractionate carbon isotopes. As a result of these uncertainties, interpreting δ13
C with
confidence can be difficult. More recent studies have attempted to quantify isotope
fractionation and the factors that affect it by growing artificial speleothems on substrates,
and then comparing the preserved isotope signature to instrumental records (Kim &
O'Neil,1997; Coplen, 2007; Tremaine et al., 2011; Feng et al., 2014).
Hendy Test
When conducting speleothem studies, it is vital to ensure that the carbonate being
studied was deposited in isotopic equilibrium. A speleothem is said to have been
deposited in isotopic equilibrium when there is equilibrium between the water and
dissolved and precipitated carbonate phases (Wong, 2015). The most commonly used
test to examine the magnitude of isotope fractionation is called the Hendy Test (Hendy,
1971). The Hendy Test involves evaluating the level of correlation between δ18
O and
δ13
C by using the coefficient of determination, R2
. If these values are correlated, then
there may be some amount of isotope fractionation that is affecting both of the signals. If
there is little or no correlation, it can be assumed that the speleothem was deposited in
isotopic equilibrium. The second part of the Hendy Test involves taking a transect of

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isotope measurements along a single lamination of the speleothem. If the isotope ratio
values are consistent along the transect, then it can be assumed that the entire speleothem
was deposited in isotopic equilibrium (Hendy, 1971). More recently, however, the
validity of this test has been called into question. It has been suggested that, because both
δ18
O and δ13
C are related to climate fluctuations, that the Hendy Test may be
fundamentally flawed (Dorale, 2009). Instead, Dorale suggested that projects with
enough funding include a second speleothem analysis from elsewhere in the same cave as
method of ensuring precision of measurements through replication. This is intuitively a
more thorough check than the Hendy Test, though it is not feasible in the case of this and
many other studies because of budget constraints and an interest in minimizing impact on
cave environments. Accordingly, this study will use the Hendy Test, and will approach
interpretations with an appropriate level of caution.
Previous Isotopic Work in The Northeastern United States
While cave speleothem studies have gained immense traction within the scientific
community over the last two decades, few caves in the Northeastern United States have
been studied in any detail. This is partially due to a lack of large cave systems, and also
due to the continuous blankets of glacial sediment that obscure the bedrock throughout
the region (van Beynen, 2004). Understanding existing paleoclimate records for this

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region will therefore be key to interpreting the Weybridge Cave speleothem isotope
record. Due to the geography of the region, most paleoclimate data for the region comes
from lacustrine sediment cores records (Kirby, 2002). This section aims to review
relevant paleoclimate information and what it suggests for this study.
In 2001, a study examined paleoclimate using δ18
O values from lake marl in a
sediment core taken at Fayetteville Green Lake, NY (Kirby et al., 2001). The core was
varved, which allowed for easy dating via a combination of layer counting and 14
C
dating. Although the record produced from this sediment core only extends to ~1000 ka,
the observed patterns provide a useful platform for identifying regional teleconnections.
The study found a 20-30 year periodicity of winter climate that persisted throughout the
entirety of the record. This sort of periodicity must be caused by an external forcing such
as solar insolation variation, or by an internal forcing such as ocean-atmosphere links.
Kirby explores both possibilities and finds that there is essentially no correlation (r = -
0.15) between modern data and solar insolation. He suggests, instead, that this regular
movement of the winter vortex can be explained by a cyclical strengthening and
weakening of the global thermohaline circulation. A weaker thermohaline circulation
system would result in less poleward heat transport, and therefore a larger winter vortex,
while a stronger thermohaline circulation would result in the opposite. It is interesting to
note that this pattern continues despite well-known, longer-duration climatic events
within their record including the Medieval Warm and the Little Ice Age (Kirby et al,
2001). More recently, Kirby published a study detailing the N-S migration of the North
Atlantic Polar Vortex in relation to consistently declining δ18
O values (Kirby, 2002).

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A frequently cited paper by Brent Yarnal (1988) uses global climate modeling and
existing instrumental data to examine decadal variations in modern Pennsylvania climate
and produces similar results. This study found that the N-S movement of winter vortex
was responsible for the major decadal shifts in precipitation quantity in the region. This
conclusion was reached by compiling past instrumental pressure records to track the N-S
movement and spatial distribution of the vortex. These findings bolster the conclusion of
Kirby’s work, and show that this pattern has continued despite recent global warming
trends (Yarnal, 1988).
More recently, Philip van Beynan (2004) conducted a study that mirrors this
project in Weybridge Cave. This project was conducted on Indian Oven Cave in eastern
New York and has produced two relevant papers; one focused on cave and regional
hydrology, and the other on the isotope record from a cave speleothem. Beynen’s
hydrology study (2006) presents a bimonthly record of regional precipitation and cave
seepage water δ18
O values. They found that Indian Oven Cave has a fairly short “flow-
through rate,” the amount of time that it takes for precipitation to enter the cave
environment. Of the two sites that they measured, one site observed a two week lag
between δ18
OPrecipitation and δ18
ODripwater, while the other site saw about a one month lag
between the two values. They determined this by offsetting and then correlating
precipitation and cave seepage δ18
O. They argue that, despite the fast flow-through rate,
paleoclimate signals are still recorded in their speleothem because the growth rate is slow
enough that each sample encompasses multiple years. Their speleothem study
encompasses a similar time frame to the one being examined in this thesis. The
speleothem was dated with U-Th series techniques and dates to 7.6 ka and showed three

23
distinct climate regimes. Based solely on their δ18
O record, they report that between 7.6-
7 ka there was a period of heightened humidity and warmth. Between 7-2.5 ka they
observed a cooling trend, and from 2.5-present they observed a relatively stable climate.
Their δ13
C values are constant, except for a period of slightly heavier δ13
C values
between 7.6-5 ka, which suggests a wetter climate, consistent with the δ18
O record.
Beynen estimates that the mean annual temperature at 7.6 ka was 4.3°C-5°C warmer than
the present.
Paleoclimate History of the Northeastern United States
Although specific knowledge of mechanisms that have controlled Holocene
climate change in the Northeast is limited, a number of studies have been conducted to
reconstruct the region’s paleoclimate (Dwyer, 1996; Mullins, 2001; Hardt, 2010; Parris,
2010). Relevant Holocene paleoclimate begins with the Younger Dryas. During the
Younger Dryas (12.5-11.3 ka), temperatures were 3-4 °C cooler on average (Schuman et
al., 2002, Zhao et al., 2010). This has primarily been determined by δ18
O values from
lake marl. Some other studies that examine pollen records from lake cores suggest that
temperatures may have been as much as 5.6 °C cooler (Yu, 2007). A brief and simplified
history of the region suggests a continuation of the cool and dry period after the Younger
Dryas between 11.6-8.2 ka, a warm and wet climate between 8.2-5.4 ka, a warm and dry
interval between 5.4-3 ka, and a cool and wet environment from 3 ka-present (Zhao et al.,
2010).
The literature generally agrees that the Younger Dryas was followed by a warm
and wet period until ~5.4 ka. This is evidenced by consistently high lake carbonate δ18
O

24
values during that time (Zhao et al., 2010). Zhao suggests that this might have been
caused by the same mechanism explored by Yarnal (1988). This mechanism suggests
that the position and retreat of the Laurentide Ice Sheet affected the position of the glacial
anticyclone circulation. This cyclone would create a northeasterly flow on the southeast
end of the ice sheet, which may explain the cool and dry climate (Zhao, 2010).
Starting at ~5.4 ka, van Beynen and Kirby’s isotope records exhibit a decreasing
δ18
O values, which correlate to about a 2° C temperature decrease. Many mechanisms
have been proposed for this change, but there is little scientific consensus on the forces
that caused this shift except that there was a decrease in solar insolation during this time.
This decrease in insolation is interpreted to broadly explain the unstable climate during
this period. This is likely because the precession index has changed solar insolation
throughout the Holocene. The precession index controls the direction of Earth’s axis tilt,
which exerts control on global seasonality. At one end of the cycle, the northern
hemisphere is closest to the sun in the summer, and at the other end of the cycle, the
northern hemisphere is closest to the sun in the winter. This controls the extent to which
summers are warm, and to which winters are cool. Modeled past changes in this cycle
show that the northern hemisphere winters, on average, have been warming, while
northern hemisphere summers have been cooling. This change in seasonal heat gradient
affects the strength of many atmospheric circulation systems, which play a large
influence in regional climate.
Lastly, most lake sediment records agree that between 5-3 ka the climate was dry
and cool (Dwyer et al., 1996; Mullins, 1998). Records from the mid-west indicate that
increased eolian sediment inputs during this time (Booth, 2005). Multiple lake studies of

25
the northeast record a decreased large precipitation storm event frequency during this
time interval as well (Noren, 2002; Parris, 2010, Munroe, 2012). This storm decrease
was recorded as a decrease in average grain size in sediment cores.
Speleothem Uranium-Thorium Dating
The data produced in this study would be difficult to interpret without an accurate
chronology tied to time of deposition. In order to make these interpretations worthwhile,
an age model linking time of deposition to calendar years before present must be
developed. Speleothems are typically dated by Uranium-Thorium dating. This technique
takes advantage of the fact that speleothems are exclusively formed through chemical
deposition by CaCO3 precipitation. They key to Uranium-Thorium dating is that
Uranium is quite soluble in water and often exists as UO2
2+
(Uranyl), while Thorium is
nearly insoluble in most groundwater. Because of this, it is safe to assume that little to no
Thorium is deposited in speleothems, while a reasonable amount of Uranium is always
deposited. Uranium concentrations can reach as high as a few hundred parts per million
(White, 2007). Because 234
Th and 230
Th are both intermediate daughters in the U-series
decay chain (Figure 5), and because it is almost safe to assume zero Th concentration at
the time of deposition, it is possible to precisely date speleothem time of deposition by
measuring the abundance of Th.
It should be noted that the amount of thorium initially present in speleothems is
not actually zero, rather, it is close to it. Some dating techniques attempt to estimate
detrital quantities of Th (Zhao, 2009), and then to subtract this amount of Th from the

26
ratio. This is done by measuring the quantity of 230
Th/232
Th in the speleothem and using
this as a proxy for the ratio of radiogenic to detrital Th (Fairchild, 2012). This process
was used in the dating of this speleothem, and the ages were corrected accordingly
(Figure 6).
Figure 5: 238
U decay series. (Fairchild and Baker 2012)

27
Figure 6: Diagram of applied detrital Th corrections based on 230
Th/232
Th ratios
The Weybridge Cave Speleothem that is the subject of this study was dated by
Zach Perzan in 2013 using this technique (Figures 5, 6, and 7). The results of this dating
show an age range of 1580 +/- 32 —4891 +/-54 years.

28
Speleothem Physical Appearance
Even when viewed with the unaided eye, it is clear that this particular speleothem
is layered, and exhibits significant color variation. The entire speleothem is
approximately 10 cm long, but the longest distance that runs perpendicular to the growth
axis is closer to 8 cm. The growth banding varies in size from 0.5mm to 5mm. The color
varies seemingly randomly from a deep brown to a light gray, and shades in-between.
This study will employ a number of techniques in order to determine the mechanism that
controlled this color change during deposition.
Figure 7: Weybridge Cave Speleothem photo with U-Th ages.

29
The speleothem stratigraphy is reasonably straightforward. It contains three
distinct growth regions with one growth region that pinches out laterally (Figure 8).
Concluding Notes
The goal of this study is to contribute to the study of paleoclimate in the
American Northeast by interpreting the geochemical data obtained from a Weybridge
Cave speleothem. Few speleothem studies have been done in this geographical region,
which provides me the unique opportunity to corroborate a small base of currently
existing knowledge about paleoclimate in the Northeastern United States. This will be
accomplished by a twofold effort. First, the speleothem itself will be drilled and analyzed
to produce continuous, high-resolution δ18
O and δ13
C records. Next, a hydrology study
will be performed on Weybridge Cave, which will involve comparing local precipitation
and cave drip water δ18
O values. This will qualitatively show how quickly water moves
through the vadose zone. These two efforts will be carefully synthesized into an
interpretation of how climate has changed in Western Vermont between 4.9-1.6 ka.

30
Methods
Speleothem Sampling Strategy
The Weybridge Cave Speleothem was sampled at a 900-µm interval in order to
produce a continuous geochemical record. The speleothem contains three distinct growth
periods, which are shown in Figure 8. Three separate transects were drilled in order to
obtain the longest distance within each region perpendicular to the growth axis (Figure
8). The sampling was done using an automated micromill at Union College. A total of
104 samples were drilled. Each sample was drilled as a 200-µm diameter hole.
Figure 8: Growth Hiatus diagram. Thick black lines represent isotope transects, while
the perpendicular, thin lines represent Hendy Test transects. The red lines represent
growth hiatuses
2
3
1
Section Analyzed by SEM

31
Scanning Electron Microscopy
In order to determine the cause of the distinct banding within the speleothem, it was
examined under the Tescan Vega 3 LMU Scanning Electron Microscope (SEM).
Because the entire speleothem did not fit into the instrument, a fragment of it was
removed (Figure 8). A region with distinct banding was chosen so that the compositional
differences could be explored with energy dispersive X-ray spectroscopy (EDS).
Water Sampling
In order to enhance my understanding of the δ18
O record, and of the mechanics by
which water is currently delivered to the speleothem location, cave drip water and
precipitation were sampled on a weekly basis for four weeks. Four collection sites were
located in Weybridge Cave, and one was located outside of the cave to collect
precipitation. Sites were chosen by looking for holes excavated in the cave floor by drip
water. Small bottles with funnels and filters were placed underneath drip sites to capture
water and to prevent evaporation. A humidity logger left in the cave recorded ~99.9%
humidity for the entirety of the sampling interval, so it is reasonable to assume that no
evaporation occurred. Sample locations were spread throughout the traversable extent of
the cave. After collection, samples were filtered, poured into small vials until full, and
then stored in a refrigerator at 5°C to prevent evaporation. The samples were analyzed at
David Gilliken’s Laboratory at Union College using ICP-MS.

32
A
B
Figure 9: Photos showing precipitation sampling methods. A: Photo of drip collector in
cave. B: Photo of above ground precipitation collector

33
Results
Age Model
Creating a robust and precise age model for this speleothem is imperative. Without
an age model, all interpretations are less significant. Creating such a model proved more
difficult than originally anticipated. To produce an accurate model, the speleothem was
examined under a microscope, and isotope drill holes were assigned to corresponding U-
Th dates. A linear growth rate was assumed between the dated intervals. Because each
sample was milled at a 0.9mm spacing interval, the remaining holes were each assigned
ages based on their linear distance from the U-Th ages. This produced the age model
shown in Figure 10.

34
Figure 10: Initial speleothem age model

35
After examining this age model, and studying the δ18
O data, it became clear that the
large decline in values around 3,900 years must be the 4.2 ka event. This event is well-
documented climate event that appears in other carbonate studies in the region (Kirby,
2002; Yu, 1997). The 4.2 ka event, or the Middle-Holocene Transition will be discussed
in detail in later sections. These other studies observe a δ18
O fluctuation of exactly the
same magnitude, therefore it is almost certain that this original age model is inaccurate.
It seems reasonable, due to the large sample size taken for U-Th dating, that the age
model could have been a few hundred years off. To refine the age model, the sample that
begins this decline in δ18
O values was assigned the age 4,200.
As the reader may note, in Figure 8 it is clear that there is an unconformity that
pinches out in the speleothem. Only about half of the unconformity was sampled and
analyzed. It is estimated that approximately 2 mm of the unconformity were not
sampled. This missing space was also added into the age model, and the growth rates
were adjusted two accommodate the two missing samples. The final age model is shown
in Figure 11.

36
Figure 11: Refined speleothem age model including Middle Holocene Transition (4.2 ka
event), and missing time.

37
Stable Isotope Data
The stable isotope data were produced using a mass spectrometer at Union college.
Figure 12: δ18
O and δ13
C Stable Isotope Data for Weybridge Cave Speleothem with
refined age model

38
As seen in the above figure, the δ18
O data seems to have two distinct means during
different time intervals, one at -7.0‰ until about 4.2 ka, and another at -7.75‰ from
3800ka to the end of the record. The gray lines labeled “Transect Break” represent
breaks in the sampling transects, which are areas of concern. In this data set, it looks as
though the samples before and after the transect breaks follow the trends that they were
previously a part of (Figures 13 and 14). Obviously some level of caution must be
observed while interpreting over transect breaks, however it may not be tremendously
important due to the consistency in trends over the breaks. It is also interesting to note
that the δ18
O and δ13
C data co-vary with an R2
value of 0.487.
Transect Correlation Coefficient of δ18
O and δ13
C
(R2
)
Whole Speleothem 0.487
1 0.340
2 0.367
3 0.384
Table 1: Correlations between δ18
O and δ13
C in different speleothem growth

39
Figures 13 and 14: Raw δ18
O values with zoomed windows. Transect breaks do not
interrupt visible trends in isotope data.

40
Figure 15: d18
O and d13
C speleothem correlation diagram

41
Hendy Test
Because this speleothem contains three distinct growth regions, it was important to
produce a Hendy Test Noncorrelation diagram for each one. As shown by the poor R2
values in Figure 16, it is likely that this speleothem was deposited in isotopic equilibrium
(Hendy, 1968).
R² = 0.11089
-7.90
-7.85
-7.80
-7.75
-7.70
-7.65
-7.60
-7.55
-7.85 -7.80 -7.75 -7.70 -7.65 -7.60 -7.55 -7.50 -7.45
d18O
d13C
Hendy Non-correlation Transect 1
R² = 0.04157
-7.44
-7.42
-7.40
-7.38
-7.36
-7.34
-7.32
-7.30
-7.28
-7.26
-7.10 -7.05 -7.00 -6.95 -6.90 -6.85 -6.80 -6.75 -6.70 -6.65
d18O
d13C

42
Figure 16: Hendy Test Non-Correlation diagrams. The lack of correlation in graph is
illustrative of the fact that this speleothem was deposited in isotopic equilibrium.
Modern δ18
O values in Vermont
In order to properly contextualize speleothem δ18
O values, it is important to
understand modern day precipitation δ18
O values. For the purposes of this study, weekly
δ18
O data from 2002–20012 was obtained from the Sleepers River Watershed in
Danville, VT USGS in Montpelier (McDonnell and Shanley). The data show a yearly
average δ18
O value of -10.8‰. The maximum-recorded value was 6.8‰, and the
minimum was -30.33‰. This gives a range of 37.13‰, which is consistent with strong
seasonality. The values in the diagram represent average δ18
O values over the last
decade for each month. The values in the summer are distinctly higher than the winter
values, and the fall and spring values typically function as a transition between the two
modes.
R² = 0.13689
-7.60
-7.55
-7.50
-7.45
-7.40
-7.35
-7.30
-7.25
-7.20
-8.40 -8.30 -8.20 -8.10 -8.00 -7.90 -7.80 -7.70 -7.60
d18O
d13C

43
Figure 17: Diagram showing the modern seasonality of δ18
O in St. Johnsbury, VT
Weybridge Cave Drip Water
Drip water was sampled in four sites within Weybridge Cave over a three-week
period in October. The δ18
O data from these samples is, unfortunately, difficult to
interpret due to the lack of δ18
O precipitation data. The precipitation catcher that was
built near the opening of the cave only successful gathered precipitation for one of the
three weeks, so it is difficult to calibrate the lag time between precipitation events, and
cave drip water. Caves with fast infiltration rates have lag time as long as two weeks, so
it is entirely possible that this data set recorded little interpretable information (van
Beynen, 2006). Nonetheless there seems to be a significant difference between sites 1-3

44
and site 4. Site 4 consistently recorded higher δ18
O values, but without knowing the δ18
O
of the input precipitation in this time period, it is difficult to speculate about the cause of
this disparity. Site 4 also has by far the largest standard deviation between samples,
which is yet another reason to resist the temptation to interpret this data. If this data is
believable, then it is possible that the drip water at site 4 is being delivered through a
different pathway, which is affecting the δ18
O values.
Figure 18: δ18
O values from 4 drip water sites inside Weybridge Cave and one
precipitation station.
-10.5
-10.0
-9.5
-9.0
Site 1 Site 2 Site 3 Site 4 Precipitation
d18O(‰)
Weybridge Cave Drip Water d18O
Variation
10/28
11/6
11/13

45
Scanning Electron Microscope Data
These images help to illustrate the cause of the changes in color within the
speleothem. Many chemical factors can contribute to the color of carbonate material,
however it is clear that subtle chemical variations are not responsible in this particular
case. From these images, it is clear that the brown layers are thin detrital concentrations
containing an array of common minerals. In addition to mapping the entire region,
individual grains were also analyzed for their composition. EDS reveals that the layers
contain plagioclase feldspar, potassium feldspar, quartz, sphene, and hematite. The major
elements were reported in percent weight, which allowed for simple stoichiometry
calculations and helped to determine these minerals. Si, Na, Fe, O, Al, Ca, Ti, Mg, S, Cl,
and K were the primary elements used to determine mineral composition. It has been
shown that Weybridge cave floods completely over time, which could explain the semi
regular presence of detrital grains (Perzan, 2014). If the cave were to fill with water, this
would halt speleothem deposition, and it is possible that small amounts of the inflow of
clay sediments could have coated the speleothem.
To further examine this assertion, the bulk chemistry of the detrital layers were
analyzed and compared against the chemistry of mud samples from the cave taken by
Zach Perzan. Data from three, combined spectrum are chemically similar to the muds
analyzed by Zach Perzan (2015), which bolsters this explanation. Fe:Ti, K:Ti, Na:Ti
ratios were examined, and all were similar.
It is somewhat concerning that the detrital layers are so thick, because this presents
the possibility that some isotope samples were centered directly on these detrital layers,
and that these detrital grains could have systematically biased the δ18
O values. To

46
evaluate this possibility, the speleothem was examined under a microscope to determine
which samples were dominated by detrital layers.

47
Figure 19: SEM photos of speleothem detrital layers at 83x and 384x zoom
Carbonate Matrix
Detrital Layer
Detrital Layer
Carbonate Matrix

48
Figure 20: EDS images showing elemental composition of detrital layers

49
Figure 21: Box plots showing the distribution of δ18
O values for detrital and non-detrital
layers in the speleothem. The difference between mean values for the two categories has
a P value of 0.05 determined with a two-tailed Mann-Whitney test (Non-Detrital n=78,
Detrital n =9).
The results of a Mann-Whitney test suggests that the difference between
samples that contained a majority detrital material, and samples that did not is statistically
significant, with a confidence interval of ~95%. This implies that the detrital layers
should not be relied on, because they artificially decreasing the δ18
O values. While it is
possible that the detrital deposition occurred only during a different climatic period which
is responsible for the difference in δ18
O values, the difference is small, and only nine
samples were impacted. In an attempt to reduce confounding variables, and to simplify

50
interpretations, these impacted data points were removed from our record.
Discussion
Modern Vermont Climate
Much research has been conducted in the Northeastern United States to
attempt to understand the atmospheric systems that deliver precipitation to the region. As
a part of the AIRMoN program, daily precipitation samples were taken in the region
since the early 1990s, many of which have been measured for δ18
O and dD. Results
show that precipitation source location exerts a first control on these values. The greatest
variation in weighted δ18
O values was observed VT99 Station (Figure 22). Further study
of this region lead to the conclusion that the δ18
O values at this northern Vermont
weather station were sensitive to precipitation source location changes (Sjostrom and
Welker, 2009). This suggests that if the relative contribution of precipitation from one
region changes in relation to another, the δ18
O values in this Vermont station are
uniquely sensitive to change.
This study identified three primary regions that contribute precipitation to Vermont
at different times of the year. There is an Arctic Precipitation source (intermediate δ18
O
values), a Gulf precipitation source (high δ18
O values), and a Pacific precipitation source
(low δ18
O values). During the winter months, the Pacific and Arctic precipitation
sources are dominant, and in the summer months, the Gulf precipitation source is
dominant. Because of this fact, the carbonate δ18
O values should be proportional to
relative amount of precipitation contributed during winter and summer. The changes in

51
source precipitation in the winter account for about 0.3‰ for every 10% change in
precipitation amount from the three contributing sources, which suggests that seasonal
precipitation source changes would have to be dramatic in order to change δ18
O values.
The change in δ18
O values seen in this record can best be explained by shifts in the
amount of precipitation contributed by summer months versus that contributed by winter
months.
Table 2: Average St. Johnsbury precipitation δ18
O values per season from 2002-2012.
Although there is some uncertainty as to modern average summer δ18
O fluctuations,
it seems clear that the relative positive and negative fluctuations over century time scales
should represent changes in precipitation source. This is because, while the amount of
summer precipitation changes, summer precipitation values are consistently more
positive than the winter values, so the δ18
O values presumably represent the relative
contribution of summer and winter.

52
Figure 22: Schematic diagram that illustrates the various modern US precipitation sources.
Stars represent studied weather stations and their names. (Sjostrom et al., 2009)
The North Atlantic Subtropical Anticyclone
One of the major controls on the quantity of summer precipitation delivered to
Vermont is the presence and strength of the North Atlantic Subtropical Anticyclone,
otherwise known as the Bermuda High. Every year in early May, this high pressure
system begins to strengthen, expand, and migrate northwards. This movement continues
until late July, when the pattern begins to reverse. This is significant because when the
anticyclone is at its peak strength, it moves warm, humid air from the Atlantic Ocean and
Gulf of Mexico to the Northeastern United States. This climate system serves as the
mechanism that delivers precipitation with higher δ18
O values to the region (Burnette
1993, Davis 1996).
Intermediate δ18
O values
Low δ18
O Values
Low δ18
O Values

53
The North American Winter-Vortex
A major control on winter precipitation is the North American Winter-Vortex, or
polar vortex, which is a high-pressure system that forms above Canada and has profound
climatic impacts on the Northeastern United States. Changes in the geometry and
intensity of this vortex have been shown to exhibit strong controls on the source
precipitation for the Northeastern United States. Historically, the southern extent of this
vortex functioned as a barrier, which deflected most oncoming Pacific precipitation.
Throughout the Holocene, the winter vortex has retracted, which has allowed for a larger
contribution of precipitation sourced from the Pacific. This northward migration should
decrease δ18
O values over time as Pacific precipitation increases.
Figure 23: Potential track of retracting North American Winter-Vortex (Kirby, 2002)

54
δ18
O Interpretation
As previously discussed, the idea of using δ18
O values as a climate proxy was
developed by Hendy in 1968. The original concept was that there is an empirical
relationship between isotopic composition and temperature of precipitation, or quantity of
precipitation. However, it has since been shown that this empirical relationship decreases
substantially when mean annual temperatures are above 10°C (Jouzel et al., 1987, 1994).
As a result, climate scientists have moved away from interpreting δ18
O values as
representing paleotemperature and past precipitation quantities. Instead, a rapidly
growing body of evidence suggests that δ18
O value fluctuations may be due to circulation
changes, or do to seasonality changes. Since the publication of Wang et al. (2001),
speleothem paleoclimate studies essentially have ceased to rely on temperature as their
explanation for changes in δ18
O in carbonate.
The trends in the δ18
O values for the Weybridge Cave speleothem exhibit a trend
towards more negative values over time, interrupted by a sharper decline around 4,200
years before present, and by a slight increase in δ18
O values around 1,500 years before
present. Moving chronologically from the oldest part of the speleothem towards the
youngest, the first notable feature is the drop from -6.5‰ to about -8‰ that occurs
around 4,200 years before present. As explained previously, it is likely that this drop
represents the impact of the Middle Holocene Transition, or the 4.2 ka event. In this
particular record, it seems likely that the 4.2 ka event is represented by the period of time
when the North American Winter-Vortex crossed Weybridge Cave on its northward path.
The Middle Holocene transition was a period of profound climate change,

55
particularly in North America. Paleoclimate studies from the Rocky Mountains indicate
that high elevation tree vegetation experienced a substantial die off, which was attributed
to a decrease in summer temperatures around this time (Benedict et al., 2008). There is a
large quantity of evidence from lake cores that suggest the occurrence of a large drought
throughout the majority of the mid-western United States starting at 4.2 ka (Dean 1993,
1997; Forman, 1995, Booth, 2004). Mid-western lakes show a sharp increase in eolian
sediment around 4.2 ka. This climatic event is present in records from every northern
hemisphere continent. There is also evidence of mass deforestation in the western
Mediterranean around 4.2 ka, which also suggests drought conditions (Magri and Parra,
2002). Evidence in North Africa suggests that many dune fields reactivated around this
time as well, presumably in response to the same drought (Swezey, 2001).
A few mechanisms have been proposed to explain the 4.2 ka event. There is no
scientific consensus as to which explanation is most probable. Proposed mechanisms
include a non-linear response to Milankovich forcing, increased volcanic activity, and
variability in the atmosphere—ocean system (Booth, 2005). There is simply not enough
high-resolution paleoclimate data from the late Holocene to know which explanation is
the most plausible. A study by Bryson in 1988 indicates an increase of volcanic activity
around 4.2 ka, however it remains unclear as to whether or not a global signal of this
increase exists in ice core records (Bryson, 1988, Booth, 2005). It has also been shown
that pacific sea surface temperature fluctuations can profoundly impact precipitation
distribution and average regional temperatures in North America. This suggests that
fluctuations of this nature, which could also decrease the strength of ENSO and create
long standing La Niña events could also be responsible. Unfortunately, until there is a

56
larger quantity of SST data for this time period, it is difficult to evaluate the relative
contribution of this forcing.
While the exact age did not originally line up with reports from the literature, the
previously explained present age model was recalibrated to properly place this event. A
similar magnitude drop in δ18
O values at this time is found in lake carbonate from
Fayetteville Green Lake in central New York (Kirby et al., 2011), and in three coeval
speleothems from West Virginia (Hardt et al., 2010). Both studies observed about 0.75-
1.0 per mil change throughout this time period. A plethora of other paleoclimate studies
record a similar δ18
O fluctuation around this time, and a most attribute this to different
combinations of the previously discussed forcings (Dwyer et al., 1996; Yu et al., 1997;
Mullins, 1998; Thompson et al., 1998; Rodbell et al., 1999; Sandweiss et al., 1999;
Claussen et al., 1999; Rodbell et al., 1999).
With the exception of the 4.2 ka event, the major driving force behind constant
decreasing δ18
O value trend between 4.2-1.5 ka is the changing solar insolation at 45°N
that has occurred through the mid Holocene. The most recent winter aphelion occurred
11,600 years ago, and Northern Hemisphere winter insolation has been slowly increasing
ever since (Berger, 1978). In the northern hemisphere, this means that average winter
temperatures have increased, while average summer temperatures have decreased. While
this change does not directly impact the δ18
O values, it drives the mechanisms that do.
The primary mechanism that contributes to this steady decrease is the constant
northward migration of the North American Winter Vortex (Kirby, 2011). The position
of the winter-vortex controls the amount of Pacific precipitation that can be delivered to
Vermont in the winter. As the vortex migrates northward, relatively depleted Pacific

57
precipitation should make up a larger percent of the total precipitation, which explains the
decrease in δ18
O value over time. Kirby asserts that the polar vortex has been moved
further north due to increasing winter solar insolation, which has caused the ice sheets to
retreat northward over time. This interpretation is consistent with my results. The
retreating winter vortex should produce a stepwise drop in δ18
O values when it moves
north of Weybridge, which is exactly what is seen.
Figure 24: Winter and Summer Insolation move divergently throughout the Holocene at
45°N

58
Figure 25: δ18
O and solar insolation roughly correlate throughout speleothem record.
Insolation data from Loutre, M.F.; Berger, A.L., 1991.
Another key component of the pronounced decrease in δ18
O values in this
speleothem record is the influence of the precession index on the summer climate system.
This is primarily reflected in the changes in the size, location, and intensity of the North
Atlantic Subtropical Anticyclone (NASA). This high-pressure system is responsible for
bringing most of the humid, Gulf moisture to Vermont during the summer, and its
intensity is directly related to summer temperatures (Hardt, 2010). Because summer
insolation has been decreasing throughout the entirety of this speleothem record, it is
reasonable to assume that the NASA has weakened over time, which would decrease
summer’s contribution to the total precipitation in Vermont. This decreasing seasonality,
combined with a greater input of Pacific moisture caused by the retracting polar vortex

59
explains the decrease δ18
O values over time.
This idea is also supported by lake level evidence. For instance, lake levels at
Crooked Pond in Southeastern Massachusetts were low at the beginning of the
speleothem record around 4,800 years BP, which is consistent with an increased summer
seasonality (Schuman and Donnely, 2006).
The clear increase in δ18
O values at about 1,700 years BP is an enigmatic deviation
from the previously described trend. A trend of this magnitude has not been produced in
other carbonate δ18
O records from the Northeastern United States, but this data clearly
represents a legitimate trend because it contains fourteen data points. Some recent paleo-
flood evidence from Vermont and New York lakes suggests that storminess peaked in the
region around 2,100 years BP, which is about as far off as this record is from the mid-
Holocene transition (Noren, 2002; Parris, 2009). It is conceivable that this increase in
storminess drove the increase in δ18
O values. Both cited studies describe the increased
summer moisture contributions as coming from the Gulf area, which would increase δ18
O
values. The majority of the records that do not record this increased δ18
O trend are from
lake carbonate, which records lake δ18
O values. The Weybridge Cave record may have a
faster response time to fluctuating δ18
O levels from precipitation because lakes should
have longer lag times. The amount of precipitation input necessary to significantly
change the δ18
O values of a large lake should be substantially larger than the amount
necessary to influence cave drip water, or carbonate δ18
O values.
It is also conceivable that this increase is simply noise in the data set, and if more
recent data existed, that δ18
O values would continue to decrease per the previous pattern.
Obtaining more modern precipitation and trace element data may help to corroborate the

60
existing data, and may help to make these interpretations more clear (Fairchild, 2009).
δ13
C Interpretation
While the initial intention of this project was to incorporate δ13
C values into my
interpretations, it has since become clear that the data produced from this speleothem are
not particularly useful. Studies that produce Holocene δ13
C almost exclusively use this
data to ensure that their carbonate was deposited in isotopic equilibrium, and usually
interpret δ13
C changes within the Holocene as noise within the data set (Kirby, 2002;
Hardt, 2010). This is mostly because the dominant vegetation in many regions did not
change during the Holocene. After considering the literature, and conferring with
experts, it has been determined that this is also the case in our record. The produced δ13
C
values can be interpreted as being influenced by a consistent and unchanging dominant
vegetation, which explains the lack of significant fluctuation in δ13
C values in our record.
Conclusion
This project produced a continuous, Quaternary paleoclimate record of the Western
Vermont region from a Weybridge Cave speleothem. The δ18
O values closely represent
what has been seen in similar studies in the region, with the exception of a noteworthy
increase in δ18
O at about 1,700 years before present. It was found that the 4.2 ka event
exerted a strong control on Vermont paleoclimate, and that δ18
O values have been
steadily decreasing since. This information helps to strengthen pre-existing knowledge
about the climate system of the Northeastern United States. It has been shown that the
retraction of the North American Winter-Vortex has profound impacts on precipitation

61
source and quantity for the region. More evidence for the speculations made by van
Beynen and Kirby about the impact of the precession cycle on this system has been
produced. The North American Subtropical Anticyclone weakens with decreasing
summer insolation, and the North American Winter-Vortex retreats with increasing
winter insolation. The combination of these factors suggests that at the current point in
the precession cycle, seasonal distribution of precipitation in this region is most evenly
distributed. At the opposite end of the precession cycle, precipitation is concentrated in
the summer, and is sourced from different locations.
The next step for this project is to refine the current age model. Due to the small
number of ages, and the relatively large pieces of material that were dated, the age model
is limited. Acquiring additional, more-focused ages for the speleothem record would
allow for more confidence in these interpretations.
The field of trace element interpretations from carbonate is also rapidly developing
(McDermott, 2003; McMillan, 2005; Fairchild, 2009; Tremaine, 2013). Future work
could include collecting a transect of trace element data from the speleothem using LA-
ICPMS. This could potentially provide insight into paleoaridity, and also bolster
confidence that this speleothem was deposited in isotopic equilibrium (Fairchild, 2009).

62
Works Cited
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70
Appendix 1: Raw δ18
O and δ13
C Data
Age (Years BP) δ
18
O δ
13
C
4891 -6.96 -6.68
4839 -6.82 -6.66
4788 -6.70 -5.90
4736 -6.88 -6.44
4684 -7.11 -7.11
4633 -7.08 -6.86
4581 -7.42 -8.08
4529 -7.39 -8.22
4477 -6.88 -6.16
4426 -7.04 -6.34
4374 -7.04 -6.55
4322 -6.88 -6.19
4271 -7.28 -7.42
4219 -6.94 -6.38
4167 -7.16 -6.00
4116 -6.84 -6.07
4064 -6.76 -6.50
4012 -6.88 -6.27
3960 -6.58 -6.25
3909 -6.92 -6.17
3857 -6.72 -6.20
3805 -7.37 -7.59
3754 -7.30 -7.16
3702 -7.77 -7.60
3650 -7.76 -6.97
3599 -7.81 -6.76
3547 -7.71 -6.26
3495 -7.64 -6.87

71
3443 -7.64 -6.92
3392 -8.11 -7.66
3340 -7.99 -8.06
3288 -7.33 -6.72
3237 -7.71 -6.70
3185 -7.42 -6.46
3133 -7.80 -7.28
3082 -7.11 -6.02
3030 -7.03 -6.07
2978 -7.20 -5.90
2926 -7.24 -6.10
2875 -7.98 -7.33
2823 -8.31 -8.22
2789 -8.47 -7.22
2756 -8.12 -8.62
2722 No Data No Data
2689 No Data No Data
2655 -7.18 -6.07
2621 -7.76 -7.52
2588 -7.54 -7.35
2554 -7.45 -7.65
2521 -7.38 -7.19
2487 -7.96 -7.15
2453 -7.38 -7.11
2420 -7.46 -6.90
2386 -7.31 -6.87
2353 -7.54 -7.19
2319 -7.56 -7.61
2285 -7.52 -6.82
2252 -7.62 -7.15
2218 -8.73 -8.81

72
2185 -8.13 -6.93
2151 -7.92 -7.68
2118 -8.05 -7.88
2084 -7.45 -8.17
2050 -7.67 -7.43
2017 -7.85 -7.80
1983 -8.11 -7.72
1950 -8.25 -9.10
1916 -7.99 -7.41
1882 -7.95 -8.46
1849 -7.95 -8.63
1815 -7.75 -8.24
1782 -7.78 -8.71
1748 -7.85 -8.25
1714 -7.95 -7.67
1681 -8.26 -9.01
1647 -7.97 -8.95
1614 -7.91 -8.52
1580 -7.71 -8.38
1546 -8.17 -7.78
1513 -7.16 -7.85
1479 -7.60 -7.96
1446 -7.88 -8.15
1412 -7.29 -7.46
1378 -7.48 -7.74
1345 -7.33 -7.87
1311 -7.53 -8.48
1278 -7.52 -7.98
1244 -7.04 -7.50
1210 -7.81 -7.31
1177 -6.93 -6.75

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