ABSTRACT
HYDROCLIMATIC VARIATIONS OF THE EARLY-TO-MID HOLOCENE IN
SOUTHWEST IRAN
By
Rosemarie H. Wrigley
August 2015
A ~2500 year record of hydrologic change from southern Iran is inferred from
the mineralogy and stable isotopic composition of bulk and biogenic carbonates
archived in Lake Hirom (27º 57’N, 53º52’E). This change is related to regional
variations in moisture and to the larger Indian Summer Monsoon circulation (ISM).
During the early Holocene, increased summer insolation from ~10,000 to 8,000 yr BP
contributed to the intensification of the ISM. This intensification may have increased
summer precipitation north of the modern ISM limit. Evidence of wetter lake
conditions in Lake Hirom occur from 8,800 to 7,800 yr BP. Drier conditions occur
and persist from 7,800 to 6,300 yr BP, until lacustrine marl changes to peat,
indicating maximum aridity. The timing of the drying trend in the mid-Holocene of
Lake Hirom correlates with cave and lake records in the Arabian Peninsula, and lake
records in northern Iran, indicating a regional drying event.
HYDROCLIMATIC VARIATIONS OF THE EARLY-TO-MID HOLOCENE IN
SOUTHWEST IRAN
A THESIS
Presented to the Department of Geological Sciences
California State University, Long Beach
In Partial Fulfillment
of the Requirements for the Degree
Master of Science in Geology
Committee Members:
Lora R. Stevens, Ph.D. (Chair)
Gregory J. Holk, Ph.D.
Matt Becker, Ph.D.
College Designee:
Robert D. Francis, Ph.D.
By Rosemarie H. Wrigley
B.S., 2012, California State University, Northridge
August 2015
ProQuest Number: 1599197
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ACKNOWLEDGEMENTS
I would like to acknowledge a number of people who have helped me with my
writing, research, and who have also given me encouragement throughout my time as a
graduate student.
I would like to express the deepest appreciation to my committee chair, Dr. Lora
Stevens who has been an amazing mentor, teacher, and advisor. I would like to thank her
for encouraging my research and for allowing me to grow as a research scientist. Her
advice on both research as well as on my career have been priceless. I am grateful for her
spirit of adventure in regard to research and scholarship, and an excitement in regard to
teaching. Without her supervision and constant help this thesis would not have been
possible. I would like to thank my committee members, Dr .Gregory Holk and Dr. Matt
Becker, for their help. Their feedback and guidance with my research and writing has
been very useful to the completion of my thesis. In addition, a thank you to Dr. Gregory
Holk and Dr. Benjamin Hagedorn whose help in the laboratory was priceless.
A Special thanks to Dr. Christine Whitcraft for allowing me to use her freeze drier
at CSU Long Beach, and to Dr. Kathleen Johnson for allowing me to use her laboratory
at UC Irvine. I would like to thank everyone in Dr. Lora Steven's laboratory for their
help with preparing and analyzing my samples. I also want to thank Gina Oliver, Tracy
La Rocco, and Priscilla Miranda for their encouragement and help with my research.
iii
I am grateful for the financial support of the Carl W. Johnson-Bert Conrey,
Graduate Fellowship in Geological Sciences. I would not have been able to complete my
degree without the financial support. My research thesis could not have been completed
without the research funds provided by NSF-EAR Grant 0903117 awarded to Dr. Lora
Stevens.
Last but not least; I would like to thank my sister, Veronica Wrigley, my mother,
Carmen Wrigley, my father, Christopher Wrigley, and my close friends Luna and Collin
Boucher. Their moral support and guidance helped me to complete my classes, research,
and thesis. There were countless times when I needed moral support, and they were
always there to encourage me and to persevere.
iv
TABLE OF CONTENTS
Page
ACKNOWLEDGEMENTS................................................................................................iii
LIST OF TABLES.............................................................................................................vii
LIST OF FIGURES............................................................................................................ix
CHAPTER
1. INTRODUCTION...................................................................................................1
2. BACKGROUND.....................................................................................................5
Site Description............................................................................................5
Regional Geology and Physiography...........................................................5
Regional Climate.........................................................................................7
Holocene Climate in Southwest Asia........................................................12
Hydroclimatic Reconstructions from Geochemical Proxies......................17
Ostracodes as a Geochemical Proxy..............................................19
3. METHODS............................................................................................................21
Methods and Proxies..................................................................................21
Core Retrieval and Sampling.....................................................................21
Chronology................................................................................................22
Ostracodes..................................................................................................22
Ostracode Geochemistry............................................................................23
Bulk Carbonate..........................................................................................25
Mineralogy.................................................................................................26
Carbon Content..........................................................................................26
4. RESULTS..............................................................................................................28
Core Lithology...........................................................................................28
Age Model/Stratigraphy.............................................................................31
Total Organic and Inorganic Carbon.........................................................32
v
CHAPTER
Page
Stable Isotopic Composition of Endogenic Carbonate...............................36
Macrofauna................................................................................................37
Stable Isotopic Composition of Ostracodes...............................................42
Trace Elements in Ostracode Carapaces....................................................44
5. DISCUSSION........................................................................................................48
Brine Evolution and Mineralogy as an Indicator of Effective Moisture....48
Detrital Minerals as Proxies of Climate.....................................................51
Isotope Data of Bulk Carbonates and Ostracodes as Indicators of
Effective Moisture.....................................................................................52
Trace Elements of Ostracodes as Indicators of Effective Moisture...........53
Mean Climate Changes..............................................................................54
Earliest Holocene (9,300-8,800 yr BP)..........................................54
Early Holocene (8,800-7,800 yr BP).............................................57
Early-to-Mid Holocene (7,800-6,300 yr BP).................................58
Mean Climate Change-Interval 1-3...............................................60
Regional Climate Change..........................................................................61
6. CONCLUSION......................................................................................................65
APPENDICES...................................................................................................................66
A. OSTRACODE OXYGEN AND CARBON ISOTOPE DATA...............................67
B. OSTRACODE TRACE ELEMENT DATA............................................................69
C. CARBONATE OXYGEN AND CARBON ISOTOPE DATA...............................71
D. TOTAL CARBON AND INORGANIC CARBON DATA....................................75
REFERENCES..................................................................................................................79
vi
LIST OF TABLES
TABLE
Page
1. Radiocarbon Dates for Lake Hirom........................................................................33
vii
LIST OF FIGURES
FIGURE
Page
1. Regional map of Asia................................................................................................4
2. Google Earth® image of Lake Hirom. ....................................................................6
3. Geologic map of southeast Fars, Iran. .....................................................................8
4. Measured precipitation in mm and calculated evapotranspiration in data over a
year...........................................................................................................................9
5. Image depicting the climatic regimes influencing southwest Iran's climate. ........10
6. SW Asia with the locations of Lake Hirom (star), lake sites (filled circles),
speleothem sites (open circles), and marine site (closed triangle). .......................14
7. Oxygen isotope record from Qunf cave, Hoti cave, and interpreted lacustrine
phases of Lake Awafi. ...........................................................................................15
8. Image of the entire 7 meters of core. .....................................................................27
9. Stratigraphic column for the entire core taken from Lake Hirom. ........................29
10. Mineral abundances of unsieved sediment by XRD. ...........................................30
11. Age–depth model of calibrated 14C dates with error bars. ...................................34
12. Total organic (TOC) and inorganic (TIC) carbon found within the marl
section of the core, grey bars indicate mean values.............................................35
13. Stable isotopic composition of the lake carbonates.. ...........................................38
14. Covariance of oxygen and carbon isotopes. ........................................................39
15. Covariance of oxygen and carbon isotopes. ........................................................40
16. Macrofuana found within the marl section of the core. .......................................41
viii
FIGURE
Page
17. The adult ostracode species count. .......................................................................43
18. Stable isotopes of the ostracod carapaces, values are reported in permil. ...........45
19. Trace elements in the ostracod carapaces, values are reported as molar ratios. ..47
20. Comparison of carbonate and ostracod carapace isotopic trends. .......................55
21. Comparison of trace element ratios Mg/Ca and Sr/Ca for C. torosa (female),
gypsum, celestine, and intermediate and high-Mg calcite (%). ............................56
22. Comparison of dry and wet climates in Lake Hirom- Iran, Lake Awafi-SE
Arabia, Hoti cave (N, north) and Qunf cave (S, south)- Oman, and Lake
Mirabad-Iran..........................................................................................................63
ix
CHAPTER 1
INTRODUCTION
The weakening and strengthening of weather systems is affected by variability in
the Earth's climate. One of the causes of the intensification in weather systems is the
global increase in temperature (Stephens and Hu, 2010). Climate model projections
predict an increase in the intensity of precipitation in wet regions and a decrease in arid
regions. Due to anthropogenic forcings we are currently experiencing global warming
and in effect we are seeing intensification of precipitation events, with strong spatial
variations (Chou and Lan, 2012). The economics of many societies are reliant on the
timing and amount of precipitation. Thus, it is important to understand how significant
these changes in precipitation are and will become, by placing these changes in the
longer-term context of Holocene climate change. Reconstructing past hydroclimatic
conditions of a region can provide links between water availability and variations in
climatic forcings, such as greenhouse gas levels, insolation, and ocean circulation.
The precession of the equinoxes, a 23,000-year cycle, changes the amount of
insolation received on Earth. Approximately 11,000 years ago the northern hemisphere
received maximum summer insolation, which caused an increase in global summer
temperatures (Ruddiman, 2000). Although not a perfect analog for today's temperature
increase, the warmer than modern temperatures during the early- to mid-Holocene. Thus,
1
this time period can be used to look at the relationship among increased annual
temperatures and effective moisture (precipitation minus evaporation).
Monsoon rains are economically vital for many countries because they are a main
source of precipitation during the summer months. Because the monsoon system
responds to the differential heating of a continent and an ocean, an increase in global
temperature and/or an increase in insolation will strengthen and potentially prolong the
monsoon. The modern Indian summer monsoon (ISM) penetrates the most southern tip
of the Arabian Peninsula near Oman in response to both the Northern Hemisphere
heating and the summer shift of the Intertropical Convergence Zone (ITCZ). Today, the
ISM does not extend past Oman and into Iran. However, the ISM was stronger in the
early Holocene (Burns et al., 2001; Parker et al., 2006; Fleitmann et al., 2007) and may
have had a greater influence in Iran.
This thesis focuses on the hydroclimatic changes during the early to mid
Holocene in SW Iran as inferred from environmental proxies archived in the lake
sediment from Lake Hirom, Iran (FIGURE 1). Lake Hirom (27º 57’N, 53º52’E) is the
most southern perennial lake site in Iran identified to date, making it potentially sensitive
to the influence of the ISM. All paleoclimatic research in Iran (e.g., Stevens et al., 2001,
2012; Griffiths et al., 2001; Djamali et al., 2009) has occurred further north beyond the
possible influence of the Indian monsoon system.
Changes in moisture balance in Lake Hirom during the early-to mid-Holocene are
identified with numerous climatic proxies, including pollen, mineralogy, sediment, and
micro- and macro fauna. Geochemical proxies are particularly useful in reconstructing
2
changes in moisture balance. The oxygen isotopic composition of endogenic carbonates
and ostracod carapaces, and the trace elements in the ostracod carapace record reflect
changes in lake water residence time and therefore variations in effective moisture. Thus,
drier times, when effective moisture is low (i.e., excessive evaporation), result in an
increase of the δ18O value of the lake water, and an increase in the relative concentration
of trace elements, both of which are ultimately recorded in the endogenic carbonate and
ostracod carapaces. These inferred changes in effective moisture can be linked to larger
changes in synoptic climate, principally the strength of the ISM and changes in the
strength of the westerly winds. This thesis will use the proxy evidence from Lake Hirom
to indicate the changes in strength of the ISM and related climatic regimes in the early-tomid Holocene.
3
FIGURE 1. Regional map of Asia. Star denotes location of Lake Hirom in SW Iran.
4
CHAPTER 2
BACKGROUND
Site Description
Lake Hirom (27º 57’N, 53º52’E), Iran is located in the southwestern region of the
Asian continent. This region is part of the Near East, which also includes countries along
the eastern shores of the Mediterranean and Northeastern Africa (FIGURE 1).
Lake Hirom lies within an intermontane basin in the southwestern portion of the
Zagros Mountains (FIGURE 2). The lake is ephemeral with large sections drying out
each year. The area with the most persistent water body is located on the northern edge
of the basin. The intermontane basin is spring fed from the base of the mountains to the
north and east (Djamali, pers. comm.). The aerial images shows that the lake can extend
~ 2 km south from the highway along the northern edge and can have an E-W width of
~2 km. The overall basin is much larger (~120km2 ) and thus during wetter intervals in
the past, the lake may have been substantially larger in area and volume. An alluvial fan
has built across the lake basin significantly limiting the size of the basin. Its role in the
water balance of the lake is unknown.
Regional Geology and Physiography
The collision of the Arabian and the Eurasian plates formed the Zagros fold and
thrust belt that trends NW- SE along the western border of Iran in the late Cretaceous.
5
FIGURE 2. Google Earth® Image of Lake Hirom. The basin is filled in the center by an
alluvial fan. The eastern edge of the basin has several small springs. The main spring is
along the northern edge of the basin (arrow indicates main spring). Black dot represents
coring site. The topographic lines were created using ASTER GDEM, which is a product
of METI and NASA.
6
The Zagros Mountains consist mainly of sedimentary rocks, with marine carbonates as
the dominant lithology (Talebian and Jackson, 2002) (FIGURE 3). Salt diapirs have been
found throughout the Zagros Mountains, with several occurring near Lake Hirom (Talbot
and Jarvis l984; Motamedi et al., 2010). The mountains form an orographic barrier
between the Mesopotamian lowlands to the west and the Iranian plateau to the east. As a
result, the Zagros Mountains have higher precipitation than the surrounding areas,
although precipitation decreases to the south (van Zeist and Bottema, 1977).
Regional Climate
Iran has a continental climate and is mainly arid or semi-arid, with an annual
temperature range of -3 to 35 degrees Celsius, with a large daily range. The Lars weather
station approximately 55 km southeast of Lake Hirom shows the annual potential
evapotranspiration and precipitation in mm (FIGURE 4). A positive moisture balance
occurs during the months of December through March, while a negative moisture balance
occurs from April through November. The wettest and driest months of the year are
December, and May/November, respectively. During July there is a slight increase in the
precipitation values.
Precipitation is controlled by the interplay of three dominant pressure systems: an
anticyclonic ridge over Asia, travelling depressions which track through the
Mediterranean Sea and the Persian Gulf, and the monsoon Asiatic low (FIGURE 5) (Taha
et al., 1981). These pressure systems strengthen and diminish according to the season,
with the first two mentioned strengthening during the winter, and the last one
strengthening during the summer.
7
FIGURE 3. Geologic map of southeast Fars, Iran. Rectangle around the Lake Hirom
Basin, star denotes Lake Hirom.
8
FIGURE 4. Measured precipitation in mm and calculated evapotranspiration in data over
a year. Most of the year has moisture deficit.
9
A
B
FIGURE 5. Image depicting the climatic regimes influencing southwest Iran's climate.
A. The winter pressure systems, the depressions (low pressure, LP) tracking with the
westerly winds and the anticylonic ridge (AR), are shown in black bold font.
B. The monsoon Asiatic low (AL) forms during the summer with the northward shift of
the ITCZ. The modern ISM extent and summer ITCZ position are shown in black bold
font.
10
The anticylconic ridge is made up by the Siberian anticyclone (northern part of
the ridge), and the subtropical high-pressure belt (southern part of the ridge, 30°N). The
Siberian anticyclone is an area of high pressure formed by an accumulation of cold air
masses (Gong and Ho, 2001). The subtropical high-pressure belt is formed by cool
sinking air from the equatorial Hadley cell (Starr, 1948). During the winter, polar
maritime air masses follow the passing Atlantic low-pressure systems through Europe,
and invade Asia Minor and the Black Sea. This leads to small depressions forming over
the Mediterranean Sea which track with the westerly winds (Kendrew, 1961). The
average frequency of the travelling depressions in the winter currently does not exceed
three a month. The extent to which these storms penetrate into the interior of Iran is
limited in part by the winter strengthening of the Siberian high-pressure system (Gong
and Ho, 2001). For most of Iran, as with the Near East in general, precipitation falls
during winter and spring (Taha et al, 1981).
Prior to the monsoon Asiatic low forming in June, the pre-monsoon circulation
consists of a split in the jets, with the polar-front jet (northern jet) and the subtropical jet
(southern jet) flowing north and south of the Tibetan plateau, respectively (Chang, 1967).
In May and June the Tibetan plateau heats up, and the subtropical jet moves south and
dissipates, leaving just the polar-front jet (Chang, 1967).
The monsoon Asiatic low forms during the summer as the heating of the Tibetan
plateau replaces the subtropical high-pressure belt. In June the Intertropical Convergence
Zone (ITCZ) shifts northward, and separates the monsoon currents from the easterlies to
the north (Chang, 1967). The convergence zone is referred to as the monsoon trough, and
11
persists until August. The western arm of the monsoon trough is the westward extension
of the low-pressure belt over India and Pakistan centered at 30°N during the summer
(Taha et al., 1981).
In the summer the warmer air temperatures allows for an increased amount of
water vapor to be stored in the atmosphere, which ultimately produces the classic
"monsoon rains"(Chang, 1967). The relative high pressure that forms over the Arabian
Sea and low pressure over the Asian continent allows for this moisture to move
northward. The modern Indian summer monsoon (ISM) penetrates the most southern tip
of the Arabian Peninsula near Oman, and does not extend into Iran.
The Shamal winds occur during the winter and summer in eastern Arabian
Peninsula and the Arabian Gulf. The winter Shamal winds are northwesterly winds that
follow the cold frontal passage over the Persian Gulf (Ali, 1992). The summer Shamal
winds are northwesterly winds that are associated with the Indian and Arabian thermal
lows, and are more likely to be dry and create dust and sand storms (Ali, 1992).
Holocene Climate in Southwest Asia
The evolution of the Holocene climate of Southwest Asia is complex and not well
understood. The early Holocene climate of the Near East (~6,000 -11,000 yr BP), in
particular, has been the focus of many paleoclimatic studies. During this time summer
insolation was greater than present with a maximum value at 523.16 W/m2 at 60°N.
Concomitantly winter insolation was lower, and the combined effect was enhanced
seasonality. The greater summer insolation caused higher summer temperatures in many
regions, and greater global climate variability. The greater summer temperatures in the
12
subtropics of the northern Hemisphere resulted in the intensification of monsoons
(Sirocko et al., 1996). Several studies indicate that the Southwest Indian summer
monsoon (ISM) intensified and may have reached more northerly latitudes in the early
Holocene (FIGURE 6) (Fleitmann et al., 2007; Overpeck et al., 1996; Burns et al., 2001;
Parker et al., 2006). In northern Iran several lake studies (FIGURE 6) have explored the
climate of the Holocene (Stevens et al., 2001, 2006, 2012; Bottema, 1986; van Zeist and
Bottema, 1977; Griffiths et al., 2001; Djamaili et al., 2009). However, these sites are
believed to be well beyond the reach of an enhanced monsoon and are interpreted in
terms of westerly airflow (Stevens et al., 2001, 2006).
In southern and northern Oman proxies indicate a rapid increase in precipitation
from approximately 10,000 yr BP to 9600 yr BP (FIGURE 7) (Fleitmann et al., 2007).
The lag between the peak solar insolation (~ 11,000 yr BP) and the peak precipitation is
attributed to glacial boundary forcings. During the deglacial at ~11,000 yr BP, pulses of
melt water weakened the North Atlantic thermohaline circulation, resulting in colder
North Atlantic air temperatures and increased snow cover in Eurasia. The increased
albedo due to Eurasian snow cover suppressed the increased summer insolation effects
which in turn delayed monsoonal precipitation until the beginning of the Holocene
(Overpeck et al., 1996; Fleitmann et al., 2007).
Sediment records from Awafi in Saudi Arabia (FIGURE 7) suggest greater
precipitation relative to evaporation beginning at 8,500 yr BP (Parker et al., 2006).
Between 12,000 and 9,000 yr BP, proxies in Awafi indicate dry conditions despite the
evidence for an intensified ISM in Oman. Parker et al. (2006) argue that the delayed
13
FIGURE 6. SW Asia with the locations of Lake Hirom (star), lake sites (filled circles),
speleothem sites (open circles), and marine site (closed triangle). Records of wetter
conditions due to the ISM are denoted with squares. The maximum summer extent of the
modern ITCZ is shown by the black line, and ISM airflow by the black arrows.
1. Marine core (Sirocko et al., 1996) 2. Southern Oman Qunff cave (Fleitmann et al.,
2007) 3. Rub a Khali lake (Parker et al., 2006) 4. Southern Oman Defore cave
(Fleitmann et al., 2007) 5. Northern Oman Hoti cave (Fleitmann et al., 2007) 6. Awafi
Lake (Parker et al., 2006) 7. Lake Mirabad (Stevens et al., 2006) 8. Lake Zeribar
(Stevens et al., 2001) 9. Lake Urmia (Stevens et al., 2012).
14
FIGURE 7. Oxygen isotope record from Qunf cave, Hoti cave, and interpreted lacustrine
phases of Lake Awafi (Parker et al., 2006; Fleitmann et al., 2007).
15
wetter conditions are due to the dominance of the Northwesterly Shamal winds, driven by
stronger westerlies. These conditions are explained by the termination of high latitude
glaciers in the early Holocene, which both weakened global wind systems and allowed
them to migrate north (Glennie and Singhvi, 2002), initiating the Shamal winds in
Southwest Asia. The increase in precipitation at 8,500 cal yr BP is attributed to the
northern penetration of the ISM (Parker et al., 2006).
Precipitation remained high in northern and southern Oman with slight decreases
at 7,900 and 7,600 yr BP. Similarly, sediment records from Awafi indicates periods of
decreased precipitation at 8,200, 7,900, and 7,600 yr BP. The weakened monsoon
activity may be due to a larger amount of snow cover in the Tibetan Plateau caused from
a melt water pulse in the North Atlantic. The higher albedo suppressed the insolation and
may have weakened the monsoon activity (Parker et al., 2006).
At 6,300 yr BP proxies in northern and southern Oman indicate a shift from
monsoon moisture to a westerly moisture source (Fleitmann et al., 2007). At ~6,000 yr
BP Awafi indicates a switch from monsoonal precipitation to cyclonic winter
precipitation from systems in the Mediterranean (Parker et al., 2006).
Northern Iran is beyond the influence of the modern ISM, however there is a
pronounced shift in hydroclimate in all records between 6000-6500 yr BP. In northern
Iran the transition to the early Holocene (~10000 yr BP), while slightly wetter than the
glacial period, is interpreted as drier or similar to present. The early Holocene evolution
is controversial with oxygen isotopes suggesting wet conditions and pollen suggesting
dry conditions. A proposed shift in the seasonality of moisture (winter precipitation as
16
opposed to spring precipitation, which would cause lower oxygen isotope values) was
used to resolve this controversy (Stevens et al., 2001, Griffiths et al, 2001). However,
new records from the Neor peatland (Arashi et al., in press) indicate lower dust
concentrations in the early Holocene relative to the late Holocene, arguing for wetter than
present conditions. These new records suggest that the oxygen isotope values in previous
studies may have indicated wetter conditions rather than a shift in seasonality that was
proposed.
Hydroclimatic Reconstructions from Geochemical Proxies
Geochemical proxies are useful in reconstructing changes in effective moisture
(i.e., precipitation minus evaporation). Oxygen isotopes of calcite from either endogenic
carbonates or ostracod carapaces can be used to indicate the effective moisture of the
region (Stevens et al., 2001; Holmes, 1996). In addition, the elemental ratios of Sr and
Mg, which substitute for Ca in the ostracod carapace, can also be used as an indicator of
salinity, and thus effective moisture as well (Chivas et al., 1986; Xia et al., 1997; Holmes,
1996; Stevens et al., 2006). These three main proxies, with additional evidence from
pollen (Djamali, pers. comm.) and mineralogy, will be used to reconstruct the
hydroclimatic evolution at Lake Hirom.
The baseline δ18O value (i.e., long-term average, prior to evaporation effects) of
the lake water is controlled by precipitation amount and type (Araguás-Araguás et al.,
2000), air mass trajectory, and temperature (Dansgaard, 1964). Seasonal timing of
precipitation (i.e., winter versus summer precipitation) can also influence the δ18O value
of lake water, where the colder temperature winter storms result in lower δ18O values
17
(Stevens et al., 2001). The trajectory of air masses feeding the lake's catchment can alter
the δ18O value of precipitation through Rayleigh distillation processes (Araguás-Araguás
et al., 2000). The further an air mass travels from the vapor source, the lower the δ18O
value of precipitation will be due to preferential rain out of the heavier 18O. Calculations
of simple Rayleigh distillation of an air mass, however, are complicated due to
entrainment of new water vapor as an air mass passes over a continental water body (ex.
Black Sea) and/or soil moisture. The monsoons create an additional complication to the
δ18O value of precipitation via a process known as the "amount effect" (Araguás-Araguás
et al., 2000). The amount effect is related to sudden and large rain out events, commonly
associated with monsoonal downpours (Rozanski et al., 1992). Thus, moisture derived
from the ISM that reaches into Iran may be strongly depleted in the δ18O due to the rain
out effect (Araguás-Araguás et al., 2000).
Once the water reaches the lake's catchment it may be further modified by
evaporative enrichment, in which the lighter 16O is preferentially removed. Evaporation
is a kinetic (non-equilibrium) process that is very difficult to quantify isotopically.
Hydrologic models of this interplay point to average temperature, relative humidity and
wind speed as critical controls on amount of evaporation (Benson, 1994). Evaporation
flux may be quantified but the isotopic enrichment due to evaporation is much more
difficult to estimate. Essentially, a comparative assessment is the only option for past
climate studies. Thus drier times, when effective moisture is low (i.e., excessive
evaporation), result in an increase of the δ18O value of the lake water, which is ultimately
recorded in ostracod carapaces and lake carbonate sediment.
18
Ostracodes as a Geochemical Proxy
Ostracodes are bivalved Crustacea that occur in either marine or terrestrial
environments. They are either benthic, living in the sediment surface or subsurface, or
periphytic, living on the surface of rooted aquatic plants. Their carapaces are constructed
with low-Mg calcite and are preserved in the sediment. There are three main lineages of
recent non-marine ostracodes belonging to the order Podcopida (Martens et al., 2008).
The largest family of this order is Cypridoedea. Two genera, Cyprididae and
Candonidae, belong to the family Cypridoedea, and are found in the paleartic zone, which
is a biogeographic ecozone that includes the terrestrial region of Europe, northern Africa,
Asia north of the Himalaya foothills, and the northern and central parts of the Arabian
Peninsula (Martens et al., 2008). Ostracodes can reproduce sexually and asexually, and
show sexual dimorphism (Butlin et al., 1998). Ostracodes, like all arthropods, molt up to
nine times, and reach lengths of approximately 0.5 to 2.5 mm long (Holmes, 1996).
Both strontium and magnesium content of simple closed-basin lakes (those with
no outflow save evaporation) can be sensitive indicator of salinity as their concentrations
increase with increasing evaporative concentration of water. Thus, the molar ratios of
Sr/Ca and Mg/Ca in non-marine ostracodes have been used in a number of studies to
reconstruct temperature and salinity changes in lakes (Chivas et al., 1986; Xia et al.,
1997; Holmes, 1996; Stevens et al., 2006). Ostracodes precipitate their shells from ions
taken directly from the water, which makes the ostracod carapace mineral composition
reflect that of the host water chemistry (Caporaletti, 2011). Because the carapaces are
secreted over a short amount of time (days), and not incrementally, their shell chemistry
19
represents the chemistry of the host water at a specific time (Holmes, 1996; Chivas et al.,
1986). During the formation of the shell, Mg and Sr will substitute for Ca in ratios
commensurate with those of the host water. Experiments have shown that the
partitioning of Sr is virtually temperature-independent (De Decker and Forester, 1988)
and is related to solely to the concentration of Sr in the host water, which typically
increases with evaporative concentration. Thus, increases in Sr/Ca ratios in ostracodes
are interpreted as increasing salinity due to a decrease of effective moisture (precipitation
minus evaporation) (Chivas et al., 1986). In contrast Mg partitioning is dependent on
both temperature and Mg/Ca of the host water (De Decker and Forester, 1988; Chivas et
al., 1986). So, changes in Mg/Ca ratios in ostracodes are more difficult to interpret. An
increase in the Mg/Ca ratio may result from decreasing water temperatures and/or an
increase in salinity. Separating the two effects is often not possible.
The partitioning coefficient for Sr and Mg is also species-dependent (Chivas et
al., 1986). Because of the molting process, juvenile ostracod shells may form during
different seasons and thus in geochemically different water than adults. Furthermore,
juvenile ostracodes preferably substitute Mg and discriminate against Sr leading to
anomalously high Mg levels and low Sr levels (Chivas et al., 1986). Therefore, adult
ostracodes are preferred for geochemical analysis.
20
CHAPTER 3
METHODS
Methods and Proxies
Proxies are used as a substitute for an indirect measurement of our past climatic
changes. In order to constrain interpretations, multiple proxies found within the marl
were investigated. The primary proxy used to infer changes in hydroclimate was the
stable-isotopic compositions of endogenic calcite and/or ostracodes. Other proxies were
used to support interpretations based on these data. All analyses, except the stableisotopic composition of ostracodes, were performed at CSULB in either the Institute for
Integrated Research in Materials, Environment and Society (IIRMES) or the Department
of Geological Sciences.
Core Retrieval and Sampling
In 2010 Morteza Djamali (IMBE-CNRS, France) collected a 6-m core at the
northern edge of Lake Hirom in approximately 15 cm of water. The core was composed
of six contiguous 1-m drives (FIGURE 8). Cores from deeper water further into the basin
were not possible without a boat, which was unavailable. The core was split in half
longitudinally, and stored in PVC pipes. Photographs were taken at 10 cm intervals for
the entire core, and a stratigraphic column was made noting the sediment structure and
type (FIGURE 9). Samples from the lower three meters of the core were collected at 5
21
cm increments for pollen analysis by M. Djamali. Pollen preservation is poor, but select
taxa are presented in this thesis. Additional samples, 5 mm in thickness, were collected
at 5 cm increments for the lower 225 cm of marl sediment. Where layering is visible,
samples were collected parallel to these layers. Contiguous samples, 5mm in thickness,
were collected in the upper 25 cm of this lower sequence to better characterize the change
in climate during the transition to the peat interval (150-375 cm). The marl samples
collected from the core were cut into two pieces, one for stable-isotope analysis and the
other for carbon content.
Chronology
A total of eight samples were collected for radiocarbon analysis (TABLE 1). The
basal peat date of 9300 cal yr BP was measured by M. Djamali at Poznań Radiocarbon
Laboratory. Six samples of plant material within the marl from 53 cm, 156 cm, 250 cm,
450 cm, 507 cm, and 600 cm and one sample within the peat from 396 cm were dated at
the Keck Radiocarbon Lab at the University of California-Irvine. All radiocarbon dates
were calibrated with Calib 7.1 (Stuiver and Reimer, 1993) and reported as cal yr BP.
Mean dates within the 2 error are used in the age-depth model. Other samples depths
were dated using a linear sedimentation rate (FIGURE 11).
Ostracodes
Sub-samples were weighed wet and dry to calculate the percent water for each
sample. This is necessary to calculate the number of ostracodes per dry g of sediment.
The subsamples were then wet sieved at 63 µm to separate ostracodes from the very fine
sediment. The ostracodes are picked from the sediment with a fine-tipped brush and
22
separated by species and by instar. The adult carapaces were cleaned with HPLC grade
ethanol for isotopic and trace element analysis, following the technique of Anadón et al.,
(2006).
Ostracode Geochemistry
Two to three adult valves of a single species were selected for δ18O and δ13C
analysis. Measurements were made on a Finnigan Mat 253 IRMS coupled to a Kiel IV
autosampler at the University of California-Irvine. Three standards, NBS-18, NBS-19,
and Oxcal, were used for calibration. The samples are placed into septa free glass vials
and phosphoric acid is reacted with the samples at elevated temperatures under
temperature control. The CO2 evolves in the septum-free vials and diffuses under
medium vacuum pressure into a crypgenic trapping system. The water and
noncondensable gases that evolved during the reaction are removed from the CO2 gas
phase under high vacuum in the first trap. If there is too much CO2 gas prior to transfer
into the microvolume, the CO2 sample size can be reduced by expansion into a defined
volume. In the microvolume, the dry CO2 is prepared for analysis in a dual microvolume
inlet system, which is connected to the IRMS. The NBS-18 standard has a δ13C value of
-5.0‰ relative to V-PDB (Vienna PeeDee Belemnite) and a δ18O value of 6.9‰ relative
to V-SMOW (Vienna Standard Mean Ocean Water). The NBS-19 standard has a δ13C
value of 2.0‰ relative to V-PDB (Vienna PeeDee Belemnite) and a δ18O value of 28.6‰
relative to V-SMOW (Vienna Standard Mean Ocean Water). The standard Oxcal, an
internal laboratory standard, has a δ13C value of -7.49‰ relative to V-PDB (Vienna
PeeDee Belemnite) and a δ18O value of 8.18‰ relative to V-SMOW (Vienna Standard
23
Mean Ocean Water). The three standards were used to normalize the raw δ13C and δ18O
values of the samples using a three point calibration method. According to the
FinniganTM Keil IV manual, the precision of this technique for δ13C is ±0.03‰ and for
δ18O is ±0.07‰. Results are reported as per mil (‰) relative to the PeeDeeBelemnnite
(VPDB) standard and are listed in Appendix A.
The phosphoric acid residue from the isotope procedure was pipetted into 15 ml
vials to which 0.5M Trace Element Grade HCl was added to keep the Sr, Mg, and Ca in
solution. This solution was analyzed on an Agilent ICP-MS at IIRMES, CSULB. An
internal standard, rhodium, is added to each sample before analysis. The sample solution
is then pumped into the inlet system from the 15ml vial. It is nebulized into a fine sample
aerosol. The aerosol is then carried into a high temperature argon plasma, which
atomizes and then ionizes the sample to produce a cloud of positively charged ions. The
sample ions are extracted from the plasma into a vacuum system containing a quadrupole
analyzer. The analyzer can scan the mass range, so multi element analysis can be
performed on the sample. The ions are focused into the analyzer, where they are
separated by their mass-to-charge ratio (m/z). The ion concentration of a specific mass-to
charge ratio is measured by an electron multiplier detector. The count rate obtained for a
particular ion is compared with a calibration plot to give the concentration for that
element in the sample. The precision of the instrument is between ±4.65% for a mass
range 7 and ±8.48% for mass range of 89. Concentrations were converted to molar ratios
and values are listed in Appendix B.
24
Bulk Carbonate
The fine sediment (<63) µm was freeze dried and crushed for bulk isotopic
analysis. Approximately 0.3 mg of sample was analyzed with a Finnigan MAT deltaplus IRMS coupled to a GasBench II carbonate device. Two international standards,
NBS-18, NBS-19, and one internal standard, STD A, were used for calibration. The
standards and samples are placed in glass vials with a rubber septa. The vials are placed
in sample trays where they are flushed with helium, which removes trapped air from the
glass tube. After the samples are purged anhydrous phosphoric acid is injected into each
vial. The liberated CO2 gas and helium enters the Gas Bench II through a diffusion trap
system, which removes the water. The CO2 is then transported to a gas chromatograph.
The isothermal gas chromatograph concentrates the CO2 gas. A reference gas injection
system is used with the sample gas for comparison. The NBS-18 standard has a δ13C
value of -5.0‰ relative to V-PDB (Vienna PeeDee Belemnite) and a δ18O value of 6.9‰
relative to V-SMOW (Vienna Standard Mean Ocean Water). The NBS-19 standard has a
δ13C value of 2.0‰ relative to V-PDB (Vienna PeeDee Belemnite) and a δ18O value of
28.6‰ relative to V-SMOW (Vienna Standard Mean Ocean Water). The Standard A, the
laboratory standard, has a δ13C value of -39.9‰ relative to V-PDB (Vienna PeeDee
Belemnite) and a δ18O value of 11.6‰ relative to V-SMOW (Vienna Standard Mean
Ocean Water). The three standards were used to normalize the raw δ13C and δ18O values
of the samples using a three point calibration method. According to the FinniganTM Gas
Bench II manual (2005 edition), the precision of this technique for δ13C is ±0.06‰ and
25
for δ18O is ±0.08‰. The Results are reported as per mil (‰) relative to VPDB and are
listed in Appendix C.
Mineralogy
The mineralogy of both the coarse (>63 µm) and fine (<63 µm) was determined
with a Rigaku Miniflex® x-ray diffractomoter (XRD) equipped with a sealed CuKα tube
and diffracted beam monochrometer to reduce β-rays. Peak identification was performed
with X’Pert HighScore Plus 3.0. The relative percentage of minerals were calculated
using Rietveld Refinement, which was performed using the structural models from the
PDF-4 minerals database with the X’Pert HighScore Plus 3.0 software.
Carbon Content
Unsieved, but crushed, sub-samples were measured for total carbon (TC) and total
inorganic carbon (TIC) analysis on a UIC™ CM5014 CO2 Coulometer. The UIC
coulometer contains a cell setup that undergoes a titration reaction when a CO2 stream is
introduced. The CO2 is quantitatively absorbed in the cell solution, monoethanolamine,
while a current generated by platinum and silver electrodes in the cell electrochemically
generates a base. The samples were weighed in ceramic and plastic boats to
approximately 11 mg. Total carbon was measured by combusting the subsample in a
furnace heated to 950° C. The total inorganic carbon was measured through acidification
of the subsample with 70% perchloric acid. The TOC is calculated by subtracting the
total inorganic carbon from the total carbon. The standard used is STD A, an internal
laboratory standard, and is ~ 12% TIC and TC, the precision of the technique to measure
26
TC is ±0.86% and TIC is ±0.80%. The TIC and TOC values are reported in percent
carbon and are listed in Appendix D.
FIGURE 8. Image of the entire 7 meters of core.
27
CHAPTER 4
RESULTS
Core Lithology
The core consists of alternating layers of peat and marl (FIGURE 8, FIGURE 9).
The basal section of the core consists of ~ 30 cm of peat that gradually transitions to
marl. The lower marl section is ~ 300 cm thick and abruptly transitions to an upper layer
of peat. The upper peat section is ~250 cm thick and abruptly transitions to the
uppermost marl section, which is ~150 cm thick. The marl sections have clear banding
with some laminated intervals. Marl sediment color ranges from rust to grey. Visual
organic matter in the marl was rare, except in the lowermost marl section.
X-ray diffractometry of the marl indicated that calcite was the dominant mineral
phase (Figure 10). The calcite exhibited varying degrees of Mg substitution. Low-Mg
calcite (<3% Mg) occurred from approximately 8,840 to 7,150 cal yr BP (630-460 cm).
Abrupt and short-lived periods of high-Mg-calcite occur after 7,150 cal yr BP, with a
higher frequency near the marl/peat transition (6,650-6,300 cal yr BP) (410-376.5 cm).
Calcite with an intermediate Mg composition occurred infrequently in the lower marl
section (8,840-7,700 cal yr BP) (630-515 cm) and increases near the marl peat transition
(6,650-6,300 cal yr BP) (410-376.5 cm). Dolomite occurred sporadically throughout the
marl section. However, it never exceeded 25% and rarely exceeded 15% of the total
28
FIGURE 9. Stratigraphic column for the entire core taken from Lake Hirom. The
locations where dates were taken are denoted with stars, age inversions are included.
29
FIGURE 10 Mineral abundances of unsieved sediment by XRD. Asterisk denotes intermediate-Mg Calcite.
Figure 10
30
mineral fraction. Periods with elevated dolomite were from 8,450 to 8,200 cal yr BP
(590-565 cm) and 7,850 to 6,300 cal yr BP (530-376.5 cm). Halite was ubiquitous with
background levels around 10%. Periods with higher concentrations occurred between
8,750 and 8,350 cal yrs BP (620-580 cm), 7,650 and 7,350 yrs (510-480 cm), and 6,550
and 6,300 years (400-376.5 cm). Gypsum was found in high percentages from 8,8407,850 cal yr BP (630-545 cm), with only very minor amounts in the uppermost marl
section. The gypsum occurred as large discoids (>63 µm) and most, if not all, is sieved
out with the macrofauna. Celestine was found throughout the core, however only in low
percentages. A single notable peak in celestine occurred at approximately 6,300 cal yr
BP (376.5cm). Quartz occurred often, with high percentages at 8,150 cal yr BP (560cm),
7,300 cal yr BP (475cm), and near the marl/peat transition at 6,550 cal yr BP and 6,300
cal yr BP (400-376.5).
Age Model/Stratigraphy
Eight dates were collected for the entire core. The marl section that is the focus
of this study was bounded by two dates on the upper and lower peat layers and three
dates on organic matter within the marl (TABLE 1). All radiocarbon ages were
calibrated with Calib 7.1 (Stuiver and Reimer, 1993). The dates at 507 cm (11,443 cal yr
BP) and 600 cm (4,665 cal yr BP) were excluded from the age model as they represent
age reversals. These two dates were excluded rather than the other dates for two reasons.
1) The peat above and below the marl is in situ. Thus, the ages retrieved using the peat
sections at 699 cm (9,300 cal yr BP) and 396 cm (6,290 cal yr BP) are the most reliable
dates. 2) The average sedimentation rate for the lower marl section, when age reversals
31
were excluded, is similar to the upper marl section (0.11 and 0.07 cm/year, respectively)
(FIGURE 11). The explanation for the age reversals at 600 cm (4,665 cal yr BP) is that it
must have been contaminated with modern carbon, and the sample from 507 cm (11,443
cal yr BP) is that it may have been re-worked. An age/depth plot (FIGURE 11) shows
the average sedimentation rates between dates and the position of the reversals. The
average sedimentation rate for the entire lower sequence is 1.06 mm/yr. Based on this
rate, each sample used in this study integrates ~5 years. The upper contiguous samples
span 24.5 cm (~125 years), whereas the rest of the sequence (lower 275 cm) have a
sample spacing of every ~50 years.
Total Organic and Inorganic Carbon
The total organic carbon (TOC) and total inorganic carbon (TIC) values of
samples retrieved from the marl are shown in FIGURE 12. TIC ranges from 0.96 to
9.92%, with a mean value of 7.98 % ( FIGURE 12). From 8,840-7,900 cal yr BP, TIC
values are generally below the long-term average and have large fluctuations. From
7,900-6,400 cal yr BP, TIC is generally above the long-term average and fluctuations are
smaller (between 7.5-10.0 %). Two notable decreases occur at 7,200 cal yr BP (470 cm)
and 6,500 cal yr BP (395 cm). In the uppermost part of the marl sequence (6,400-6,330
cal yr BP), TIC values vary significantly (0-10%). In general the TIC values increase
from the early to mid-Holocene, then decrease at the marl/peat transition. TOC ranges
from 0 to 9.3%, with a mean value of 1 % ( FIGURE 12). From 8,840-6,400 cal yr BP
TOC values are low (0-2 %), with frequent variations. Near the top of the sequence in
the mid-Holocene (6,400-6,330 cal yr BP), the TOC values show large fluctuations up to
32
TABLE 1. Radiocarbon Dates for Lake Hirom
14
C Years
Depth
Sample
Before
Age (cal yr
(cm)
Lab ID
ID
present
error BP)
type
53
UCIAMS
152321
1420
60
1336
Charcoal
156
UCIAMS
103156
2110
104
2090
Charcoal
250
UCIAMS
103157
3305
95
3525
Charcoal
396
UCIAMS
103158
5480
28
6290
Charcoal
450
UCIAMS
152322
5955
20
6767
Charcoal
507
UCIAMS
152323
9985
30
11443
Charcoal
600
UCIAMS
152324
4100
20
4665
Charcoal
699
POZ
31054
8290
50
9300
Charcoal
33
FIGURE 11. Age–depth model of calibrated 14C dates with error bars. The ages at 507
cm (11,443 cal yr BP) and 600 cm (4,100 cal yr BP) were excluded from the model
because they were interpreted as age reversals.
34
FIGURE 12. Total organic (TOC) and inorganic (TIC) carbon found within the marl
section of the core, grey bars indicate mean values.
35
10%. The TOC variations do not show a significant trend, other than an increase in the
mid-Holocene.
Stable Isotopic Composition of Endogenic Carbonate
The carbon and oxygen isotopic composition of endogenic calcite is shown in
FIGURE 13. It should be noted that hi-Mg calcite and low-Mg calcite were not
separated, and it is expected that the isotopic fractionation will differ between these two
(Tarutani et al., 1969; Land, 1980).
The small amounts of dolomite were also not taken into account. No corrections
can be made for the dolomite given that it is likely detrital and rock samples were not
available. However, there is no coherent offset between the isotopic values of samples
with or without dolomite (FIGURE 15).
δ18O values range from -4.6‰ to 0.25‰ with a mean of -2.5‰ and a standard
deviation of ± 0.9. From 8,840 to 7,800 cal yr BP, the δ18O values become isotopically
lighter over time with the lowest values at -4.6‰ at 7,800 cal yr BP. There are distinct
positive peaks during the general negative trend that occur at 8,700 and 8,050 cal yr BP.
From 7,800-6,300 cal yr BP, δ18O values range from -4‰ to -1.5‰, and vary around the
average. δ13C values range from -7.9‰ to -1.5‰ with a mean of -4.22‰ , and a
standard deviation of ±1.16. From 8,840 to 7,800 cal yr BP, the δ13C values are generally
below the average value, with a distinct and protracted decrease at -7.9‰. There are also
positive peaks during the general negative trend that occur at 8,700, 8,200 and 8,050 cal
yr BP. The δ13C values increase to -1.5‰ at 7,700 cal yr BP showing greater short term
variability, with a range from -2‰ to -7‰.
36
Macrofauna
The macrofauna found in the lower marl section consist of gastropods,
foraminifera, and ostracodes (FIGURE 16). Gastropods are abundant in the top section
of the marl (greater than 20 per interval) (7,800-6,300 cal yr BP) (525-376.5 cm), but less
abundant in the lower section of the marl (1-5 per interval) (8,840-8,550 cal yr BP) (630600 cm). They are absent between 8,500 and 7,800 cal yr BP (595-530 cm). A single
species of foraminifera is found in abundance in the lower-most, middle, and top most
section of the marl (greater than 20 per interval) (8,840-8,800, 7,600-7,550, and 6,6506,300 cal yr BP) (630-625, 510-500, 410-376.5 cm). The species is tentatively identified
as Ammonia tepida, a saline, non-marine benthic species found in Egypt and Israel
(Hayward et al., 2004).
Ostracodes do not occur in the lower-most section of the core, unlike gastropods.
They appear around 7,800 cal yr BP and increase in abundance to 6,300 cal yr BP (525376.5 cm) (FIGURE 17). Ostracodes show poor diversity and are represented by two
main species, Cyprideis torosa (Jones, 1850) and Candona candida (O.F. Muller, 1776)
with rare occurrences of Eucypris inflata (Sars, 1903). Identifications of C. candida and
E. inflata are tentative. Cyprideis torosa exhibits clear sexual dimorphism, so the female
and male species were grouped separately. Females were more abundant than males in
many intervals. The tentatively identified Candona candida was found in several
intervals where Cyprideis torosa did not occur. E. Inflata occurs only twice in the marl
section in the upper-most section of the early Holocene segment, though they are the
dominant species in the youngest marl sections above the peat (Stevens, pers. Comm.).
37
FIGURE 13. Stable isotopic composition of the lake carbonates. The grey bar denotes
the average values for the carbon and oxygen isotopes.
38
F
FIGURE 14. Covariance of oxygen and carbon isotopes. The lack of a pattern between
low- and high-Mg calcite indicates that fractionation has not caused significant changes
in isotopic values.
39
FIGURE 15. Covariance of oxygen and carbon isotopes. The lack of pattern for dolomite
indicates that fractionation did not cause a significant change in the isotopic values.
40
FIGURE 16. Macrofuana found within the marl section of the core. Circles indicate the
presence of the macrofauna within the specific interval.
41
Stable Isotopic Composition of Ostracodes
Carbon and oxygen isotopic values of ostracod carapaces are shown in FIGURE
18. The entire ostracode record represents ~160 years, thus changes are considered very
abrupt and short-lived. Limited data are due to the lack of ostracodes in the lower marl
section, thus the ostracode record is significantly smaller than the sediment record.
Analyses were done on individual species and, in the case of Cyprideis torosa, by gender.
The data are not continuous, providing only windows of information. Data often do not
overlap, but where they do usually do not track well between the different ostracodes.
The δ18O values of Candona candida have an average value of 0.47‰. The δ18O values
of Cypredies torosa male have an average value of -0.86‰, with much greater variability
than C. candida. Peaks occur at 6,510, and 6,450-6,440 cal yr BP (395-393, and 390-387
cm). A pronounced increase occurs from 6,430 to 6,400 cal yr BP (387-384 cm). The
δ18O values of C.torosa female have an average value of -1.83‰. From 6,480 to 6,430
cal yr BP (392-387 cm) the δ18O values are lower than average. They return to their
average value by 6,430 cal yr BP and are roughly constant with only small variations.
The carbon isotope pattern is quite different. As with the 18O values, there is offset of
about 3 to 5 ‰ between the C. candida and C. torosa ostracodes. However, unlike the
18O values, the torosa male and female 13C values track very well. The δ13C values of
Candona candida vary between about -2 and -6‰, but there are too few data for a profile
interpretation. The δ13C values of Cypredies torosa male have an average value of 7.1‰, with a standard deviation of ± 2.0‰. The δ13C values of Cypredies torosa female
have an average value of -6.60‰, with a standard deviation of ± 3.20‰. The most
42
FIGURE 17. The adult ostracode species count.
43
marked feature of the record is a decrease in δ13C values for both C. torosa between
6,520 and 6,490 cal yr BP (395 and 393cm). Values are steady around -8 ‰ until 6,410
cal yr BP (385 cm), and then decrease again. At 6,400 cal yr BP (384 cm) the carbon
isotopes increase and remain around -8‰.
Trace Elements in Ostracode Carapaces
The molar ratios for Mg/Ca and Sr/Ca in the ostracodes are shown in FIGURE 19.
The molar ratios are multiplied by103 to better depict the changes in the values. It is not
expected that the different species or gender will have the same ratios. The Mg/Ca ratios
for both ostracode species generally have the same trend. Candona candida has an
average Mg/Ca value of 9.7. Cyprideis torosa (male) has an average Mg/Ca value of 9.5.
Cyprideis torosa (female) has an average Mg/Ca value of 10.33, with a standard
deviation of ±2.47. The Mg/Ca values for both species stay around the average from
7,600 to 6,390 cal yr BP (510-383cm). An increase in the Mg/Ca values occurs from
6,400 to 6,330 cal yr BP (384-379 cm). The Sr/Ca values for both species are variable.
Candona candida has an average Sr/Ca value of 16.39 that generally increases over time.
Cyprideis torosa (male) has an average Sr/Ca value of 22.32. The Sr/Ca values for
Cyprideis torosa male fluctuate around the average value, although at first the values
begin slightly lower from 6,500 to 6,490 cal yr BP (394-393cm). Cyprideis torosa
(female) has an average Sr/Ca value of 22.90. The Sr/Ca values for Cyprideis torosa
female fluctuate around the average value, except at 6,370 to 6,360 cal yr BP (381380cm), where there is a large increase.
44
FIGURE 18. Stable isotopes of the ostracod carapaces, values are reported in permil.
45
Mg/CaWater and Sr/CaWater molar ratios were derived using a partitioning
coefficient value for the species of ostracod Cyprideis from De Decker and Forester,
1988. Cyprideis torosa (male) indicates the host water had an Mg/Ca molar ratio of 2.16
with a standard deviation of ±0.82, and Sr/Ca molar ratio of 0.05 with a standard
deviation of ±0.008. Cyprideis torosa (female) indicates the host water had an Mg/Ca
molar ratio of 2.25 with a standard deviation of ±0.54, and Sr/Ca molar ratio of 0.05 with
a standard deviation of ±0.011.
46
FIGURE 19. Trace elements in the ostracod carapaces, values are reported as molar
ratios.
47
CHAPTER 5
DISCUSSION
Changes in hydroclimate can cause lakes to change geochemically and physically.
These changes affect the sediment, macrofuana, and vegetation within the lake. Changes
in the lithology and faunal assemblages can be used as paleoclimate proxies themselves,
but also archives additional proxies. Both lithology and macrofauna suggest a lake of
moderate salinity. Thus, multiple lines of evidence can be used to infer climatic changes.
Brine Evolution and Mineralogy as an Indicator of Effective Moisture
There are three dominant types of sediment found along the edge of Lake Hirom:
peat, marl, and evaporites. Today, vegetation is concentrated at ground water seeps along
the edge of the lake, and thus the peat likely indicates periods with lower lake levels and
a water table just below ground level (Djamali, Pers comm.). Lake Hirom appears to be a
closed basin as seen in FIGURE 2, which suggests that drainage did not control peat
development, rather changes in lake level did. Ultimately peat at the coring site, as
opposed to marl or evaporites, suggests a lack of open water and thus drier conditions.
The first transition of marl to peat is at 6300 cal BP, thus indicating the onset of drier
conditions in the mid-Holocene relative to the early Holocene.
The sediment of Lake Hirom shows significant mineralogical changes that
suggest some type of brine evolution. Much of the sediment record is dominated by marl
48
with evaporite minerals occurring several times mixed in with the marl. The marl is often
low-Mg calcite but can include high-Mg calcite and even minor amounts of dolomite.
The presence of high-Mg calcite implies an evolution of the lake water to more saline
conditions. Evaporation of a lake closed to surface water outflow, like Lake Hirom, will
lead to a relative enrichment of solutes. The sequence of minerals that precipitate out of
the water depends on the composition of the water entering the lake and saturation of
different minerals (Hardie and Eugster, 1970). There are no current data available on the
water chemistry of either the lake or the springs feeding the lake. So conclusions
regarding brine evolution are only tentative.
The dominant carbonate phase in the lake is low-Mg calcite (<3 molar %). Both
the calcite and bicarbonate are derived from the surrounding mountains, which have a
dominant carbonate lithology (FIGURE 3). The precipitation of intermediate-Mg
calcite* (~3molar %) and high-Mg calcite (> 5 molar %) occurs when Ca removal via
calcite precipitation is significant (Müller et al., 1972). The removal of Ca causes an
increase in the Mg/Ca ratio, driving the increased Mg substitution into the calcite lattice.
Increases in Mg/Ca ratio will continue to occur with an increase in the residence time of
the lake water and lack of renewed Ca input. Thus, intermediate and high-Mg calcite are
interpreted as drier conditions (Müller et al., 1972).
As the residence time increases, bicarbonate may also be exhausted causing the
precipitation of gypsum (Hardie and Eugster, 1970). Should the Ca used in gypsum
precipitation be consumed, it may be possible to precipitate (or co-precipitate) another
sulfate mineral or even halite. Celestine, SrSO4, may precipitate once Ca is consumed
49
(McCarten et al., 1988). Halite is able to precipitate once little Ca remains in the brine as
well (Müller et al., 1972).
The lower-most part of the marl contains low and intermediate-Mg calcite,
gypsum, and halite. Though the halite is ubiquitous throughout the core it does generally
increase when gypsum increases in the lower-most part of the marl. The upper part of the
core has intermediate and high-Mg calcite, celestine, and halite. There are no
relationships found between the minerals in the upper part of the core, so no clear
explanation of brine evolution can be proposed.
The frequent occurrence of halite suggests that there is a constant supply of
sodium and chloride to the lake. Likely there is a salt dome nearby—perhaps below
ground level. Salt domes are common in the Zagros Mountains (Motamedi et al., 2011).
Thus halite is likely not a useful proxy of effective moisture indicator for this region.
The presence of peat suggests prolonged dry conditions. As opposed to marl,
which indicates wetter conditions. The sporadic occurrence of gypsum in the lower part
of the core suggests large and abrupt droughts, which caused the lake to have an
excessive brine evolution. The mineralogy of the lower part of the core suggest a
transition from prolonged droughts (peat) to a wetter climate that was punctuated by large
but short duration droughts (marl and evaporites). The mineralogy in the upper part of
the core suggest increased aridity until peat occurs, which indicates prolonged dry
conditions.
50
Detrital Minerals as Proxies of Climate
In Lake Hirom, quartz and dolomite are considered to be detrital in origin. The
quartz, which is more resistant to physical and chemical breakdown may have been
transported long distances (Dearing, 1999). Quartz is found frequently throughout the
core, but rarely occurs in high percentages. Given that most of the quartz is less than 63
µm, it is entirely possible that it is windblown and related to dust from the Shamal.
However, relative increases in percentage of quartz occurs with increases in 18O values,
which suggest that the dust is greater, with greater aeolian deposition, during drier
conditions, although the source maybe somewhat local. Drier conditions would
destabilize vegetation and wind would entrain the quartz. The co-occurrence of quartz
with celestine confirms that quartz is likely indicative of drier windier conditions.
The occurrence of dolomite is problematic. Dolomite in lake sediment can be a primary
precipitate (Warren, 1999), diagenetic (Mees et al., 2011), or detrital (Sabins, 1962).
Precipitation of dolomite directly from water requires special conditions and excessively
high Mg/Ca ratio (Müller et al., 1972). The fact that the lake precipitates hi-Mg calcite
indicates that this could be possible—but the hi-Mg calcite does not correlate with the
dolomite. Furthermore if the dolomite were endogenic it would be precipitated in larger
quantities during periods interpreted as drier, which it was not. Thus, dolomite is not
believed to be a primary precipitate. Diagenetic alteration of calcite to dolomite is
common in sabhkas (Kendall, 1979). The alteration of calcite to dolomite would cause
an increase in Ca in the pore water, and therefore a possible increase in diagenetic
Gypsum precipitation. However, dolomite and gypsum do not increase concurrently at
51
Lake Hirom. Thus, dolomite is considered to be detrital, which is plausible as it is less
soluble than low-Mg calcite. Dolomite, derived from the Mesopotamian lowlands, is a
common constituent of windblown dust into the Arabian Sea (Cullen et al. 2000). As
with quartz, dolomite may be derived from long distances and its source is unknown.
The quartz and dolomite mineral percentages do not have a relationship suggesting a
pattern with detrital mineral input. The detrital minerals are therefore just a common
background pulse of dust and unhelpful in establishing drier conditions
Isotope Data of Bulk Carbonates and Ostracodes as Indicators of Effective Moisture
The δ18O and δ13C values of certain hydrologically closed basins have been
shown to have a positive covariance (Talbot, 1990; Li and Ku, 1997), although the
relationship is equivocal. Evaporation in closed basins results in an increase in δ18O
values. Simultaneously the uptake of 12C via algae causes an increase in δ13C values of
DIC from which the carbonates precipitate. Interestingly the carbonate carbon and
oxygen isotopes in Lake Hirom do not show a strong covariance (FIGURE 14). The lack
of covariance can be a result of three factors. The first is that the lake is hydrologically
open, which is physically not obvious (FIGURE 2). The second is that different 18O
values contributed from the fraction of dolomite and high-Mg calcite disrupted the
covariance. The third is that covariance is not a robust indicator of closed basins. If for
example, carbonate precipitation is triggered by degassing rather than photosynthetic
uptake of CO2 by algae (Kelts and Hsu, 1978), then there is no a prior reason, particularly
in a low productivity lake, that the two would increase simultaneously.
52
Bulk sediment carbonates are often avoided for stable isotopes analysis in
paleoclimate studies because it may be difficult to distinguish between authigenic and
allogenic carbonates (Leng and Marshall, 2004). Because the mountains that surround
Lake Hirom are limestone dominated mountains (Talebian and Jackson, 2002), there is a
possibility that some of the carbonates used for isotope analysis were detrital. Because
the ostracodes use the water to construct their carapaces, endogenic carbonate should
mirror the isotopic trend of the ostracodes. The given caveat is that there are strong vital
effects in the fractionation of isotopes into ostracodes (Xia et al., 1997). In FIGURE 20
we see that the isotopic values of the ostracodes and carbonates follow similar trends,
within reason. They will not track each other perfectly as the carbonates may precipitate
in different seasons than the ostracodes precipitate their shells. But larger trends should
follow and they do. Thus, it is inferred that the isotopic signature from the bulk
carbonates are derived predominantly from endogenic carbonates.
Trace Elements of Ostracodes as Indicators of Effective Moisture
The concentration of Sr and Mg of the lake water can be used to infer the
moisture balance, because excess evaporation will decrease lake volume, which causes an
increase in the relative concentration of various solutes. While excess precipitation will
increase lake volume, which causes a decrease in the relative concentration of various
solutes in the water. An increase in Sr and Mg will occur concomitantly with an increase
in celestine, and/or intermediate-and high Mg calcite. A comparison of Sr/Ca and
celestine shows that the Sr/Ca ratios do increase as the celestine occurrence increases
(Figure 21). The Mg/Ca ratios of the lake water extrapolated from the ostracodes were
53
approximately 2.2, and values above 2 should result in high-Mg calcite precipitating
(Mueller, 1972).
There is no relationship between increased Mg/Ca ratios and high-Mg calcite.
However, there is a relationship between increased intermediate-Mg calcite and increased
Mg/Ca ratios. The co occurrence of increased Sr/Ca and Mg/Ca ratios and celestine and
intermediate -Mg calcite, respectively, suggest that these proxies are a good indicator of
effective moisture.
Mean Climate Changes
The early Holocene lacustrine history is divided into three intervals, which are
defined by changes in the sedimentology, mineralogy, and general trend of the isotopes.
Names are given to each interval for discussion purposes only. These names are not
formal designations. The three intervals are Interval 1: 9300-8800 yr BP (Earliest
Holocene), Interval 2: 8800-7800 yr BP (Early Holocene), and Interval 3: 7800-6300 yr
BP (Early-to-mid Holocene).
Earliest Holocene (9,300-8,800 yr BP)
The earliest Holocene (Interval 1) is composed of peat, which is interpreted as
drier than present conditions. Today, vegetation is concentrated at ground water seeps
along the edge of the lake, and thus the peat indicates either extremely low lake levels or
a water table below surface (Djamali, Pers comm.). Both support generally drier
conditions. The gradual change from peat to marl indicates that the transition to wetter
conditions was not abrupt.
54
FIGURE 20. Comparison of carbonate and ostracod carapace isotopic trends. The
oxygen trends for both proxies correlate indicting that carbonate is largely authigenic.
55
Mg/Ca and Sr/Ca for C. torosa (female), gypsum,
celestine, and intermediate and high-Mg calcite
(%). Grey bars show regions where Mg/Ca and
Sr/Ca concentrations have increased.
56
FIGURE 21 Comparison of trace element ratios of Mg/Ca and Sr/Ca for C. torosa (female), gypsum, celestine, and
intermediate and high-Mg calcite(%). Grey bars show regions where Mg/Ca and Sr/Ca concentrations have increased.
Figure 221 Comparison of trace element ratios
Early Holocene (8,800-7,800 yr BP)
The sediment transition from peat to marl is interpreted as a transition to a
permanent or semi-permanent water body due to wetter conditions. The sediment is
dominated by low-Mg calcite with punctuated occurrences of gypsum. There are some
foraminifera and gastropods in the lowest part of Interval 2, but there is a distinct lack of
macrofossils , for reasons that are unclear. So isotopic values are restricted to those from
calcite. The lack of high-Mg calcite and infrequent dolomite suggests that isotopic shifts
due to mineralogic changes are minor. A long-term trend to more negative 18O values
during this interval suggests that the overall climate was getting wetter. However, six
abrupt drought events superimposed on this trend are inferred from increased 18O values
(8,700, and 8,050 cal yr BP), decreased TIC% (8,700, 8,500, 8,400, and 8,050 cal yr BP),
and peaks in gypsum % (8,700, 8,500, 8,400, 8,200, 8,050, and 7,950 cal yr BP). The
latter suggests draw down of the lake for multiple years with the resultant increase in
18O values.
The disappearance of gastropods and foraminifera occurring between 8,800 and
7,800 cal yr BP is concurrent with several drought events in Interval 2. Drought events
may cause lake desiccation, which is coupled with a decrease in macrofauna diversity
and/or removal of less adapted species (Hunt and Jones, 1972). The fact that macrofauna
do not occur frequently after the first drought event can be interpreted several ways. 1)
The macrofauna were not able to recover after the drought events. However, if the
macrofauna were able to recover after a drought, as they did in the first drought event,
they were not preserved during the subsequent lake desiccation events. 2) Macrofauna
57
prefer to live in submergent vegetation (Hunt and Jones, 1972), and therefore the overall
wetter trend of the lake may have caused the lake to grow in size, and any submergent
vegetation had followed the shoreline away from the coring site.
Early-to-Mid Holocene (7,800-6,300 yr BP)
The transition between Intervals 2 and 3 is not obvious in the stratigraphy but is
defined by changes in mineralogy. At the beginning of Interval 2 the disappearance of
gypsum suggests an end or decrease in the drought events. However, the occasional
occurrence of celestine may signify shorter drought events as water becomes more
concentrated with Sr. Celestine forms in evaporate-carbonate environments when reacted
with Sr-rich fluids (Hanor, 2000), but it is not clear why Sr would be the dominant cation
over Ca, save for the fact that Mg is also increasing in the calcite at this time.
In Interval 3 TIC% and TOC% fluctuations around the mean. Like the
mineralogy the TIC% and TOC% have large fluctuations near the end of Interval 3,
although the rapid timing of these changes is likely due to the contiguous sampling
interval. The decrease in TIC% indicates a transition from calcite to a more concentrated
brine mineral, celestine in this case. The increase in the TOC% indicates an increase in
vegetation, which is corroborated by an increase in the aquatic pollen found in this
section of the core (Djamali, Pers comm.). The 13C values in Interval 3 are more
positive relative to Interval 2. The increase in the 13C values also indicate an increase in
vegetative productivity.
The 18O values of the bulk carbonate in this interval are more positive relative to
Interval 2. The 18O values fluctuate around the average 18O value of -2.5‰, and this is
58
interpreted as the lake fluctuating between drier and wetter stages. Increased 18O values
do not correlate with increased intermediate- and- hi-Mg calcite percentages, which
indicate that fractionation between low and high-Mg calcite does not cause a large
difference in the 18O values (FIGURE 14). However, it is somewhat perplexing if both
are driven by evaporative concentration.
The appearance of macrofauna at the beginning of Interval 3 suggests that either a
suitable substrate developed at this time, allowing for habitation and preservation, or that
climatic/aquatic conditions became more suitable. Macrofauna are primarily affected by
the salinity of lakes (Geddes et al., 1981), but their distribution can also be affected by
organic matter content and substrate (Rieradevall and Roca, 1995). Rieradevall and Roca
(1995) documented that with increasing salinity there is a decrease in the diversity of
macrofuana. However the diversity of ostracodes in Lake Hirom is already quite poor.
Salinity does not seem to influence the diversity of macrofauna as seen by the general
increase in the benthic ostracodes and gastropods occurring concurrently with a change in
mineralogy, a decrease in TIC%, and an increase in trace elements. This indicates that
submergent vegetation and a stable substrate is the more likely driver in presence/absence
of macrofuana in this lake.
The dominant ostracod, Cyprideis torosa, is a euryhaline species and does not
offer much information concerning salinity (Heip, 1976). However, both the female and
male forms are useful for isotopic and trace element evaluations of water quality. From
6,510 to 6,440 cal yr BP the decrease in the 18O value of C. torosa indicates a shift
toward wetter conditions. The 18O values increase after 6,440 cal yr BP to 6,330 cal yr
59
BP to 6,330 cal yr BP indicating that the lake is getting more concentrated and the
climate drier.
The 13C values C. torosa can be used to indicate lake vegetative productivity,
and in turn effective moisture. From 6,510 to 6,480 cal yr BP the decreasing 13C values
indicates a decrease in lake productivity. After 6,480 cal yr BP the 13C values increase
until 6,420 cal yr BP indicating increased lake productivity. The 13C values slightly
decrease again from 6,420 cal yr BP to 6,330 cal yr BP, indicating a slight decrease in
lake productivity. The changes in productivity indicate a sporadic occurrence of aquatic
plants, which suggests a transition to more peat like conditions.
High Sr/Ca values for C.torosa occur concomitantly with the increase of celestine
around 6,400-6,330 cal yr BP, which is consistent with concentration of Sr in the water
relative to Ca. Similarly the increase in high-Mg calcite and moderate-Mg calcite near
the end of Interval 3 coincides with an increase in C. torosa Mg/Ca values as well. The
presence of increased Mg/Ca and Sr/Ca values, celestine, and intermediate and high-Mg
calcite suggest that there is excess evaporation, indicating drier conditions.
Mean Climate Change-Interval 1-3
The occurrence of peat in the earliest Holocene (Interval 1) suggest that the lake
had a prolonged dry climate. In the early Holocene (Interval 2) the lack of macrofossils
and increased evaporites indicate short duration, but large scale drought events.
Although, the overall oxygen isotope signature in Interval 2 indicates increasingly wetter
conditions. The early-to-mid Holocene (Interval 3) was marked by an increase in lake
water concentration as seen by the mineralogy and ostracod Sr/Ca and Mg/Ca ratios.
60
Also, an increase in aquatic plants and TOC% indicates a transition to peat. The proxies
in Interval 3 suggest that the Lake Hirom is getting drier as it approaches the midHolocene.
Regional Climate Change
Southwest Asia (Iran, Turkey, Iraq, Saudia Arabia) is unique because it is on the
nexus of several synoptic climate systems. The interplay of these systems determines
when and how much precipitation the region will receive. A major question concerns
whether SW Iran was dominated by the ISM, or by westerly rainfall during the early
Holocene. A comparison of the results of Lake Hirom with other climatic records can
help elucidate this question (FIGURE 22).
During the early Holocene (8000-10000 yr BP), increased summer insolation and
decreased winter insolation caused greater seasonality. Lake and speleothem records in
the Arabian Peninsula suggest that the ISM reached more northerly locations as early as
10,000 yr BP (Overpeck et al., 1996; Burns et al., 2001; Fleitmann et al., 2007; Parker et
al., 2006). However, it is unlikely that monsoonal rains would have penetrated as far
north as Hirom, given the occurrence of peat and overall dry conditions. The strength of
the ISM is inferred to have increased gradually from 10500-9500 cal yr BP in Oman
(Fleitmann et al., 2007). But the prevailing dry conditions from 9300-8800 yr BP at Lake
Hirom and the similarly dry conditions at Awafi suggest that if the ISM did strengthen,
its effects were not dramatic at either location. In Oman the monsoon strength remained
high until ~6300 yr BP, where northern Oman began to see a decrease in precipitation
from ISM. In contrast Awafi had a dry climate at 9000 cal yr BP and a change to a
61
wetter climate at 8500 cal yr BP, suggesting that if monsoonal rains did penetrate this far
north it was until well into the Holocene (Parker et al., 2006). Alternatively, Awafi could
be fed by westerly precipitation, although Parker et al. do not support this model. At
~6000 yr BP the ISM weakens, but Awafi is still has marl mixed with sand after ~6000 yr
BP and only intensifies to present (Parker et al., 2006). Thus, this region of Saudia
Arabia became more arid at 6000 yr BP at nearly the same time as Hirom.
Lake Hirom has a similar oxygen isotope trend to Lake Mirabad in northern Iran.
Given the much more northerly position of Mirabad, the trend was hypothesized to
indicate an increase in winter storms (snow) rather than an increase in overall
precipitation. This leads to two possibilities to explain the lower 18O values in Interval 2
(the early Holocene).: either the ISM penetrated as far north as Lake Hirom or increased
winter storms occurred as far south as Lake Hirom.
At 8,500 cal yr BP both Awafi and Lake Hirom developed a permanent water
body, which is consistent with either the northern penetration of the ISM or the
southward penetration of the westerlies. Parker et al. (2006) favor the former, although
no concrete evidence is provided to exclude the latter. Given that these two climatic
systems operate during different seasons, it is possible that they are not mutually
exclusive. If the ISM penetrated to Hirom, then the descending arm could be over
northern Iran (dry and possibly early summers). This would mean that the majority of
precipitation in Mirabad occurred as winter rain/snow, as suggested by Stevens et al.,
(2006) and Griffiths et al., (2001). However, a dust record form Lake Neor (37°57'37"
N, 48°33'19" E), northern Iran suggests that during the early Holocene this region was
62
FIGURE 22. Comparison of dry and wet climates in Lake Hirom- Iran, Lake Awafi-SE
Arabia, Hoti cave (N, north) and Qunf cave (S, south)- Oman, and Lake Mirabad-Iran.
The asterisks indicate drought events, the likely source of precipitation is written above
each site.
63
indeed wet, not dry (Al Sharifi et al., in press). This would suggest that the shamal were
less active and that the westerlies were stronger.
Wet winters do not preclude summer rain at Lake Hirom, but do not support it
either. If the model of a northward displaced ISM is valid, then its constancy can be
examined by the isotopic record. Drought events, or precipitation minima, are considered
to be times when the ISM is weakened. With the tentative chronology, five significant
droughts 8,500, 8,400, 8,200, and 7,950 cal yr BP events occurred at Hirom. Two of
these 8,200, and 7,950 cal yr BP are in agreement with droughts documented in Awafi.
The 8,200 cal yr BP event is also correlative with a precipitation minima at Oman,
suggesting a link between the ISM and all three sites. However, it also corresponds to
Bond Event 5, which is a cooling event from a melt water pulse in the North Atlantic,
which would have impacted all synoptic climate patterns (Bond et al., 1997).
A cessation of speleothem formation in Oman at 6.3 ka BP signals the retreat of
the ISM to the south. This effectively moved the ITCZ and may have impacted all
synoptic systems in the region. Awafi, Hirom, and Mirabad all experience a fundamental
climatic shift at this time—to a drier mean climate state. Not all are linked to the ISM, so
whether the drying at Hirom is due to loss of summer monsoon rains or a large scale shift
in westerlies still cannot be resolved.
64
CHAPTER 6
CONCLUSION
This thesis adds to the collection of records that document climatic changes in
Southwest Asia. The Lake Hirom proxy record includes carbon content, sediment type,
minerology, trace elements of ostracod carapaces, and stable isotopes of ostracod
carapaces and carbonates. The earliest Holocene (9,300 cal yr BP) is marked by peat
formation and indicative of dry conditions at a time when the ISM was increasing in
strength in north and south Oman. The dry conditions in the earliest Holocene coincide
with dry conditions in Awafi Lake in south east Arabia. The stable isotope record from
8,800 to 7,800 cal yr BP has decreasing δ18O values, which indicates wetter conditions.
The timing of the wetter conditions is in good correlation with a permanent water body
forming in Awafi, a site interpreted as being influenced by the strengthened ISM
penetrated (Parker et al., 2006). One drought event in Lake Hirom at 8,200 cal yr BP has
been recorded in all the lake and speleothem records in the Arabian peninsula and
corresponds to Bond Event 5 (Overpeck et al., 1996; Burns et al., 2001; Fleitmann et al.,
2007; Parker et al., 2006). The formation of peat in Lake Hirom at 6,300 cal yr BP
indicates a transition into prolonged dry conditions. The timing of the drier climate
coincides with drier conditions occurring in northern lake sites in Iran, and in Awafi.
65
APPENDICES
66
APPENDIX A
OSTRACODE OXYGEN AND CARBON ISOTOPE DATA
67
Depth
δ13C
cm
C.candida
379
379.5
-5.3703
380
380.5
-5.0912
381
-3.1588
383.5
384
384.5
385
386
386.5
387
-4.9284
387.5
-6.2400
388
-6.0956
389
-2.3745
390
392
393
394
395
480
505
δ18O
0.2813
-0.7924
-0.2630
-0.6020
-0.7865
-1.0558
-0.0551
δ13C
C.torosa (m)
δ18O
-7.2715 2.0471
-9.6003 -1.2391
-8.6609 0.8576
-10.2475
-9.2858
-7.3726
-7.4070
-6.4099
3.1917
-0.3577
-1.5980
-1.6748
-2.9420
-8.2073
-8.6408
-7.2007
-7.3474
-4.0937
-0.2680
-0.4570
-2.8941
-9.4257 -2.8941
-6.1169 2.9699
-3.7256 -0.3920
68
δ13C
C.torosa (f)
δ18O
-8.5972
-9.0553
-8.0056
-9.6287
-11.4802
-9.3541
-0.0287
-0.7436
-0.1434
-1.4036
-0.7177
0.2912
-7.5010
-7.8438
-6.6283
-0.0666
-1.4525
-1.5342
-8.1229
-9.1664
-8.9420
-4.4266
-5.5423
-3.9478
-8.5901
-10.2390
-3.0015
-2.9154
0.0129
-2.2724
-0.7679
-1.2467
-5.6854
-2.7392
APPENDIX B
OSTRACODE TRACE ELEMENT DATA
69
Depth
cm
Mg/CaOstracod
x103
Sr/CaOstracod
x103
C.candida
Mg/CaOstracod
x103
Sr/CaOstracod
x103
C.torosa (m)
Mg/CaOstracod
x103
Sr/CaOstracod
x103
C.torosa (f)
379
19.4903
19.4903
16.8622
21.2503
9.9605
20.9786
379.5
6.7181
19.4903
16.8622
21.2503
8.7815
22.4782
2.8196
22.6033
11.4692
30.6594
7.7577
23.8133
9.7793
32.1006
13.8607
31.3694
6.4466
23.0029
8.9902
19.5922
380
380.5
7.3186
19.2174
381
8.3479
20.2182
383.5
384
6.7768
20.3948
384.5
10.1986
30.3912
385
9.6361
20.8262
7.9315
22.8011
386
13.0643
23.1999
12.9030
17.8580
386.5
12.0738
26.1915
11.8921
24.6785
387
8.8768
13.7874
387.5
9.4574
13.7113
11.3155
26.7013
11.6029
25.3941
388
10.7371
12.6731
8.7461
14.5262
9.4072
26.6901
389
6.5629
12.4958
6.6383
23.5603
7.6459
25.2276
8.9327
21.9978
7.4213
19.5171
390
392
393
7.2709
21.9978
8.7319
13.8956
394
10.3261
16.1100
14.8135
22.1929
14.1709
16.0327
10.1253
17.7908
395
480
505
70
APPENDIX C
CARBONATE OXYGEN AND CARBON ISOTOPE DATA
71
depth δ13Ccarbonate δ18Ocarbonate(VPD)
cm
‰
‰
376.5
-4.51
-2.48
377
-4.67
-2.53
377.5
-4.25
-2.87
378
-4.20
-2.60
378.5
379
-3.87
-2.53
379.5
-3.89
-2.52
380
-3.89
-2.52
380.5
-3.85
-2.23
381
-3.65
-3.08
381.5
382
-4.46
-1.75
382.5
-4.09
-1.11
383
-4.00
-1.54
383.5
-4.06
-2.41
384
-2.89
-2.69
384.5
-2.44
-2.47
385
-2.58
-2.48
385.5
-2.69
-3.02
386
-3.71
-3.40
386.5
-4.26
-3.30
387
-4.02
-3.61
387.5
-3.98
-3.36
388
-4.61
-4.08
388.5
-4.58
-3.65
389
-6.59
-3.92
389.5
-3.91
-3.01
390
-4.16
-2.37
390.5
-3.89
-2.64
391
-4.26
-2.98
391.5
-4.05
-2.51
392
-3.86
-2.58
392.5
-3.09
-1.86
393
-3.26
-2.36
72
depth δ13Ccarbonate δ18Ocarbonate(VPD)
cm
‰
‰
393.5
-3.34
-1.50
394
-3.64
-1.65
394.5
-3.54
-1.12
395
-3.60
-1.16
395.5
-3.61
-1.44
396
-2.79
-0.49
396.5
-2.95
-1.17
397
-4.38
-1.45
397.5
-3.96
-2.68
398
398.5
-3.38
-2.60
399
-4.06
-2.56
399.5
-3.77
-2.97
400
-4.13
-2.57
400.5
-4.48
-2.13
401
-4.07
-2.27
405
-3.58
-2.61
410
-2.26
-3.58
415
-4.50
-1.57
420
-4.09
-3.55
425
-2.74
-3.00
430
-3.97
-1.14
435
-4.01
-2.90
440
-6.78
-3.26
445
-4.00
-2.34
450
-4.40
-3.04
455
-3.69
-2.23
460
-3.58
-2.39
465
-5.71
-3.38
470
-4.63
-1.30
475
-4.27
-1.43
480
-3.69
-2.81
485
-4.62
-1.81
490
-5.64
-2.27
495
-5.30
-3.32
500
-3.81
-2.99
73
depth δ13Ccarbonate δ18Ocarbonate(VPD)
cm
‰
‰
505
-3.78
-2.84
510
515
-1.61
-2.50
520
-3.76
-4.64
525
530
-7.10
-3.92
535
-7.04
-4.13
540
-6.89
-3.61
545
550
-4.16
-1.55
555
-6.29
-3.48
560
-7.87
-2.60
565
-2.64
-3.23
570
-4.14
-2.44
575
-4.73
-3.21
580
585
590
595
600
605
610
-6.22
-0.41
615
-5.03
0.22
620
-6.47
-1.05
625
630
-5.60
-2.45
74
APPENDIX D
TOTAL CARBON AND INORGANIC CARBON DATA
75
depth
cm
376.5
377
377.5
378
378.5
379
379.5
380
380.5
381
381.5
382
382.5
383
383.5
384
384.5
385
385.5
386
386.5
387
387.5
388
388.5
389
389.5
390
390.5
391
391.5
392
392.5
393
TIC
%
7.70
7.98
3.79
8.63
2.30
9.35
8.03
8.40
0.01
8.78
8.83
9.21
8.13
8.28
8.61
8.42
7.75
8.45
9.24
9.58
8.27
9.75
9.14
9.59
9.63
9.81
9.52
9.43
9.93
9.44
9.33
8.57
6.51
6.25
TOC
%
0.00
0.00
1.70
0.00
6.71
0.00
1.17
0.00
9.31
0.00
1.89
0.96
1.65
2.22
1.28
1.15
2.35
1.33
0.52
0.53
1.96
0.86
1.19
0.00
0.00
0.87
1.32
1.01
0.65
1.33
0.82
1.34
2.00
1.46
76
depth
cm
393.5
394
394.5
395
395.5
396
396.5
397
397.5
398
398.5
399
399.5
400
400.5
401
405
410
415
420
425
430
435
440
445
450
455
460
465
470
475
480
485
490
495
500
TIC
%
7.19
7.34
7.84
8.08
8.48
7.96
8.56
8.54
9.23
9.39
9.59
9.45
9.24
9.28
9.59
9.10
8.33
9.53
9.16
9.36
8.38
8.67
9.00
8.87
7.84
9.60
9.38
8.79
8.87
5.29
8.36
8.71
8.04
8.92
8.77
7.58
TOC
%
0.00
0.93
1.65
0.58
1.28
2.44
0.47
1.86
0.33
0.21
0.10
0.64
0.97
0.00
0.00
0.74
1.53
0.00
0.00
0.89
0.98
0.00
1.93
1.02
0.00
0.29
0.66
0.67
0.40
0.34
1.40
0.39
0.81
0.98
1.11
0.00
77
depth
cm
505
510
515
520
525
530
535
540
545
550
555
560
565
570
575
580
585
590
595
600
605
610
615
620
625
630
TIC
%
8.56
8.62
7.81
9.58
7.78
9.07
8.05
6.83
9.75
2.84
4.88
6.39
8.10
7.21
9.20
6.83
3.85
8.97
5.41
8.04
8.10
3.21
0.96
3.42
7.22
8.02
TOC
%
1.45
0.09
0.50
0.99
0.84
0.00
0.00
1.03
0.56
0.00
0.08
0.72
0.26
0.70
2.08
2.26
0.58
1.03
1.10
0.21
0.68
0.47
0.15
0.93
1.26
1.26
78
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