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Description of global C-cycle
omparison of different time scales
On geological time scales, global atmospheric CO2 is regulated by the slow (millions of years) C-exchange between the atmosphere, ocean, buried organic matter and lithosphere (Berner, 1994; Berner and Kothavala, 2001). The main driving factors are volcanism (bringing C to the atmosphere) and continental weathering associated with plate tectonics (removing C from the atmosphere). On shorter time scales, the C-cycle is dominated by faster processes of atmosphere-ocean gas exchange, as well as by the growth and decay of the terrestrial biosphere, through photosynthesis and respiration. Fig. 1-2 illustrates this difference.
The natural carbon cycle relevant to Quaternary time scales
In the framework of this study, C cycle will be examined from a CO2-centric point of view, since, foremost, the CO2 evolution is of main interest for this study and secondly, because CO2 acts as a “conveyor” of C-exchange between two major reservoirs, the terrestrial biosphere (including the soils) and the ocean (including sediments). The amount of CO2 in the atmosphere is small, when compared with the gross exchange fluxes between the three main reservoirs. The ocean and terrestrial biosphere amounts to 38000 and 2000 GtC3, respectively while the atmosphere represents only ~730 GtC. Atmospheric CO2 is therefore expected to react sensitively to changes in the global C cycle.
Before the industrial revolution of the mid 1800’s, CO2 was stable and the C-reservoir exchanges were under a dynamic equilibrium. We actually recognize the perturbation of this dynamic balance by the human intervention (cf. I.1), further seen in the right panel of fig. 1-2 (asredarrows andnumbers)
• Terrestrial Biosphere
CO2 exchange between the atmosphere and the biosphere results from three different processes: photosynthesis, respiration and biodegradation. These processes are directed by the biochemical reaction, described by the two-way equation (1-1): Plant photosynthesis corresponds to the conversion of water and CO2, under the influence of light and specific enzymes, to glucose molecules and oxygen (right direction of eq. 1-1). Respiration is the opposite reaction, where the plants consume oxygen, rejecting CO2 in the atmosphere. When respiration occurs in the soil, it is called biodegradation, and atmospheric CO2 is produced by the oxidized (“decomposed”) organic matter.
Some numbers
Total CO2 quantity dissolved in the leaves stomata amounts to ~270GtC/y, i.e. 1/3 of total atmospheric CO2 (Farquhar et al., 1993; Ciais et al., 1997). The “fixed” quantity, consists of CO2 converted to carbohydrate during photosynthesis and corresponds to the Gross Primary Production (GPP). Ciais et al., 1997 measured it through the stable oxygen isotopes of atm. CO2 and found it to be of ~120 GtC/y. It is equally supposed that half is incorporated in new tissues (leaves, roots, wood) and the other half is converted to CO2 with autotrophic respiration (Lloyd and Farquhar, 1996). Net Primary Production (NPP) corresponds to the difference between photosynthesis and respiration; it is estimated to be of ~60GtC/y (IPCC, 2001). NPP returns to atm. CO2 pool via two processes: heterotrophic respiration, Rh and combustion. Fig. 1-2 gives provides more numbers.
• Ocean
The ocean covers 3/4 of the Earth’ surface, thus constituting a major area for C exchange with the atmosphere. It acts either as a source or a sink for CO2 through various physical, chemical (both related to temperature, salinity, pH..), dynamical (e.g. circulation) or biological processes.
Physical-chemical processes
In the ocean, CO2 exists in three different inorganic forms: as free CO2 (~1%), CO2,aq , as bicarbonate ion, HCO3 (~91%) and as carbonate ion, CO32 (~8%), all three constituting the Dissolved Inorganic Carbon, DIC. A fourth form is the true carbonic acid, H 2CO3 , whose concentration is much smaller than that of CO2,aq (≤ 0.3%) (Zeebe and Wolf-Gladrow, 2003). Surface water temperature changes influence the atmospheric CO2 value, resulting from Henry’s law of temperature-dependent gas dissolution effect on liquids (cooler water enhanced CO2 dissolution lower atm. CO2). CO2 dissolution in water also depends on salinity and alkalinity (pH indirectly). The equilibrium between gas and dissolved phase is represented in eq. 1-2: CO2, gas ¬¾®CO2,aq. (1-2) K0 Where K 0 the solubility coefficient of CO2 in seawater.
Then, dissolved CO2 reacts with water, to form carbonic acid, H2CO3, a weak acid leading to a double dissociation:
K1 + H (1-3)
H 2O + CO2 « H 2CO3 ¬¾® HCO3
K 2 2 + H (1-4)
HCO3 ¬¾®CO3
K1 and K 2 are the equilibrium constants and depend on the temperature and salinity, the latter having the opposite effect on solubility than temperature.
The partitioning between the three inorganic forms of CO2 also depends on the pH, or the alkalinity of waters (Sanyal et al., 1995). Alkalinity (Alk) corresponds to the water capacity in capturing a sufficient number of H+ protons, which permits the complete dissolution of carbonic acid (eq. 1-3 and 1-4). The transformation from gaseous to dissolved CO2 is favoured by rising the pH (reducing the water acidity), since the pH is related to carbonates (CaCO3) chemistry. Therefore, a pH rise transfers atmospheric CO2 into the surface ocean.
Thermohaline circulation and CO2 cycle
Temperature or salinity changes in the surface water at the air/ocean interface provoke changes in density. Such modifications can generate an instability of the water column, accompanied by a new regime of water flow, following the new density configuration. Ocean circulation driven by the interplay of temperature and salinity gradients, is defined as thermohaline circulation (THC). Deep cold and salty waters are formed at high latitudes of both hemispheres: in the N. Atlantic, more specifically the Labrador, Norwegian and Greenland seas for the NH, and the S. Ocean and the Weddell Sea for the SH. They are represented as NADW5 and AABW6 in fig. 1-4. The former move southwards, where they join the latter and together they raise (upwell) to the surface in the Pacific/Indian regions. This circuit is closed back to the Atlantic ocean, by a surface counter-current, principally driven by Coriolis force and winds. This oceanic journey largely contributes to climate patterns throughout the globe (Rahmstorf, 2002).
In principle, enhanced deep water formation implies reduced atm. CO2 values due to stronger ocean pumping. Still, reality is more complex since deep water formation enhancement in one ocean area may provoke increased upwelling elsewhere, thus equilibrating the CO2 oceanic sources and sinks.
Biological processes
The marine “biological pump” first consists of photosynthesis, i.e. CO2 assimilation from microscopic algae (phytoplankton) under light presence, producing living organic matter (either dissolved or particulate, DOC or POC7, respectively). Apart from soft tissues, marine biological processes also produce shells, containing carbonate (CaCO3) particles in surface waters; this is called “(counter) carbonate pump”. IPCC, 2001 reported 90 GtC/y of exchanged CO2 through this step, from which 88Gt are returned to the atmosphere.
The phytoplankton is consumed in a second step by the zooplankton (planktonic animals), which is further eaten by bigger predators, (e.g. fish). Therefore, the initially photosynthesized carbon is either transformed in short-living detritus stock, such as skeletons or faecal particles or is already being demineralised by the respiratory pathway. If the latter occurs in surface waters, CO2 is directly exchanged with the atmosphere; the percentage of this reaches ~90%, according to IPCC, 1995. A non-negligible quantity is furthermore exported in a third step towards the deep ocean, either through organisms effectuating vertical transects or under detritus form being denser than water (e.g. dead phytoplankton cells, animal skeletons in form of calcareous shells). 45 GtC/y is the calculated flux for these transports. A very small fraction of this (of the order of 0.01 GtC/y) will be definitely stored in marine sediments during millions of years and will later constitute the sedimentary rocks, which on geological time-scales will return as CO2 in the atmosphere through volcanism (cf. left panel of fig. 1-2).
Biological processes require light, CO2, nutrients (nitrogen, phosphorous) and oligo-elements (Fe). Nutrients in particular are known to be the main limiting factor for most of oceanic regions. An exception, more discussed in Ch. IV , concerns the HNLC8 regions: such regions (e.g. S. Ocean, N. and Equat. Pacific) experience low productivity, as deduced from the low chlorophyll content, while their nutrients are abundant. They are therefore supposed to be biologically limited by iron, Fe (Martin et al., 1990; Kumar et al., 1995).
CO2 evolution in different past time scales
Beyond ice cores
A visualization of the actual atmospheric CO2 behaviour is already presented in the left panel of fig. 1-1, from the beginning of continuous direct atm. CO2 measurements in 1958, until now (mid-2008). In the right panel, the temperature anomaly is illustrated. For both cases, similarities are distinguished.
The notion of global warming is linked to the Earth’s climate. If one compares this situation with previous decades, or even hundreds of thousands of years, neither temperature (via proxies), nor CO2 show such big values. Still, on time scales of millions of years (Ma or My), the atmospheric CO2 mixing ratio was far more elevated than nowadays, reaching 20-fold higher CO2 values than present, during the early Phanerozoic (until ~550 Ma). Fig. 1-6 illustrates this difference in atm. CO2, in a cover up of the whole Phanerozoic, compared to model predictions for the future. Atmospheric CO2 in such early periods has been estimated from equation-based CO2 variations governing the outgassing and weathering reactions, formulated in the GEOCARB model of Berner, 1994; Berner and Kothavala, 2001.
Fig. 1-7 leans on the Cenozoic era (until 60 My) and the relation of CO2 with ice volume (expressed through δ18O of benthic foraminifera). The so-called “hyperthermals”, occurring during the Eocene and Paleocene, are warm events accompanied by elevated CO2 concentrations (e.g. Mid-Eocene Climatic Optimum, Early Eocene Climatic Optimum). Based on these observations, a number of studies suggest CO2 to be the driver for these warm phenomena (Zachos et al., 2008; Bains et al., 2000; Berner and Kothavala, 2001; Bowen et al., 2004; Lowenstein and Demicco, 2006), following the baseline theory first proposed by Arrhenius, 1896 and which possibly has extensions to nowadays. The opposite effect is observed for later (closer to today) time scales, such as the Mid-Miocene Climatic Optimum, an important transition in the Earth’s history, marking the passage from Greenhouse to Icehouse climate status. During this optimum, temperature is elevated but CO2 values remain low (Pagani et al., 1999; Holbourn et al., 2005). For this second case, theories lean on the leading role of tectonics or orbital forcing on the temperature configuration.
For such early time scales, CO2 is reconstructed from:
• boron measurements in the ocean
• δ13C from alkenones
• paleosols (carbonate nodules in ancient soils)
• stomata density of fossilized leaves
These methods are applicable to all geologic periods, the Cenozoic being a period of major interest. Nevertheless, they suffer from high uncertainties. Ice cores provide the oldest direct means in reconstructing the atm. CO2 with lower uncertainties than the methods exposed before.
Ice core studies
The oldest up-to-date direct record of atm. CO2
Overall, it seems difficult to deduce an unifying historical relationship between CO2 and climate by extracting a simple correlation between the two variables throughout the Earth’s history, as other important factors (notably plate tectonics affecting atmospheric and oceanic circulations and thus climatic patterns) intervene. It thus appears appropriate to focus on more recent time periods, when plate tectonics did not change significantly the configuration of Earth surface.
High-resolved ice core records cover the recent Quaternary period. They permit to examine the relative phasing between CO2 and temperature, thus extracting useful information on leads and lags for the last 800 ky BP9. Fig. 1-8 presents the correlation between CO2 and temperature, the latter being reconstructed from the deuterium/hydrogen measurements. The tight relationship between these two variables is obvious, since warm events are accompanied by elevated CO2 concentrations and vice versa.
Glacial- interglacial cycles
The recent 800-ky CO2 record, corresponds to 8 climatic cycles. These cycles are represented by saw-tooth shaped curves involving a rapid temperature rise followed by a slow decline (see framed box). Similar behaviour is observed for a number of climatic variables, originating from various means (marine sediments, continental flora, fauna and loess) other than ice cores. These records reveal a main 100-ky glacial (G) /interglacial (IG) periodicity, accompanied by secondary frequencies of 41, 23 and 19 ky, the socalled orbital parameters (eccentricity, obliquity and precession of the Earth’s orbit). The leading role of these “Milankovitch cycles” on the G-IG transition has been widely recognized.
Definitions
The transition from glacial to interglacial state (“deglaciation”) is widely expressed by the “Termination, T” notion (Broecker and van_Donk, 1970), followed by a Latin number (I, II, III etc.). Further back one goes in time, higher the number is. In ice core data, Ts are best represented by: (i) the deuterium/hydrogen isotopic record, δD, an atmospheric temperature indicator showing higher values for IG state and (ii) the atmospheric oxygen isotopic record, δ18O2, a “proxy” of mainly global ice volume and, consequently, sea level changes, which gets depleted during IG state (Broecker and Henderson, 1998).
Table of contents :
CHAPTER I. INTRODUCTION
I.1. General context
I.2. The global carbon cycle
I.2.1. Coupling between C-cycle – climate: greenhouse effect
I.2.2. Description of global C-cycle
I.3. CO2 evolution in different past time scales
I.3.1. Beyond ice cores
I.3.2. Ice core studies
I.4. Stable carbon isotopes – a climatic tool
I.4.1. Introduction on the notion of isotopes
I.4.2. δ notion – application on C – δ13C definition
I.4.3. Isotopic fractionation
I.4.4. Previous studies using carbon isotopes to constrain the global C cycle
I.5. CO2 and δ13CO2: combined reconstruction from ice cores
I.5.1. Significance in measuring δ13CO2 in ice cores
I.5.2. State of the art
I.6. Scope of this study
CHAPTER II. THE ICE CORE ARCHIVE
II.1. Introduction
II.1.1. Interest in studying polar ice cores
II.1.2. Ice cores of interest during this PhD
II.2. Gas trapping in the ice
II.2.1. Densification process
II.2.2. Gas transport within the firn column
II.3. Physical processes affecting gases occluded in ice
II.3.1. Gravitational separation of gases and isotopes
II.3.2. Thermal diffusion
II.3.3. Diffusion along a concentration gradient
II.3.4. Bubble close-off processes
II.3.5. Clathrate formation and decomposition
II.4. Chemical processes affecting gas composition in ice
II.4.1. Greenland observations
II.4.2. In situ reactions of organic compounds (mainly oxidations)
II.4.3. In situ reactions of inorganic carbonate, CaCO3, followed by reactions under acidic environment
II.5. Extension of physical processes on gas age calculation challenge
II.6. Conclusions
CHAPTER III. METHODS AND INSTRUMENTATION
III.1. Introduction
III.1.1. Analytical challenges related to quantified isotopic measurements
III.2. Overview of gas extraction methods on ice samples
III.2.1. Wet extraction or melt-refreezing method
III.2.2. Dry extraction method
III.2.3. Extraction by sublimation
III.2.4. Intercomparison of the analytical system between 3 European laboratories (EPICA Consortium)
III.3. General concept and layout of the experimental procedure
III.3.1. Overall
III.3.2. sample preparation
III.3.3. air extraction
III.3.4. gas transfer
III.3.5. analysis of (i) CO2; (ii) its stable isotopes
III.4. Description of the analytical procedure
III.4.1. preparative steps before analysis
III.4.2. Strictly time-dependent procedure
III.4.3. Protocol changes through time (2006 vs. 2007-8)
III.5. Corrections applied
III.5.1. Corrections by the IRMS software
III.5.2. Other corrections
III.6. Procedure validation with blank tests and reference gas standards
III.6.1. Laboratory calendar
III.6.2. Results from standard gases
III.6.3. Results from blank measurements
III.7. Analytical observations on ice core measurements
III.7.1. CO2 and isotopic ratio trends with expansion
III.7.2. Reproducibility on numerous replicates of Vostok ice core
III.7.3. Extraction efficiency – bubble against clathrate ice
III.8. Conclusion
CHAPTER IV. TERMINATION I
IV.1. Introduction
IV.2. State of the art
IV.2.1. Reasons for studying this transition in EDC core
IV.2.2. Rapid climate changes presentation
IV.2.3. Paleoceanographic proxies used to interpret the data
IV.3. CO2 and δ13CO2 results over TI from the EDC core
IV.3.1. Laboratory brief calendar
IV.3.2. Main draft
IV.3.3. Supplementary material
IV.4. Discussion – On the causes of CO2 deglacial evolution
IV.4.1. Sea Surface Temperature
IV.4.2. Sea level
IV.4.3. Sea ice extent
IV.4.4. Ocean circulation
IV.4.5. Wind speed
IV.4.6. Coral reef hypothesis
IV.4.7. Carbonate compensation
IV.4.8. Rain ratio change
IV.4.9. Fe fertilization hypothesis
IV.4.10. Role of terrestrial biosphere
IV.4.11. Synopsis
IV.5. Additional “young” points to our dataset
IV.6. Synchronization issues
IV.7. Comparison with unpublished δ13CO2 data
IV.7.1. UBern
IV.7.2. AWI
IV.8. Conclusion
CHAPTER V. TERMINATION II
V.1. Introduction
V.2. State of the art
V.2.1. Why studying TII period on EDC ice core
V.2.2. Main abrupt climatic changes during TII
V.2.3. Absence of proxies from N.H. ice cores
V.3. Results
V.3.1. Laboratory brief calendar
V.3.2. Data processing
V.3.3. CO2 and δ13CO2 time evolution
V.4. Comparisons with other TII data in ice
V.4.1. Deuterium
V.4.2. Vostok CO2
V.4.3. CH4
V.4.4. Deuterium excess
V.4.5. nss-Ca+2 and Fe fluxes in EDC core
V.4.6. ss-Na+ flux in EDC core
V.4.7. paleo-oceanographic data over TII
V.5. Comparison between TI and TII
V.5.1. Amplitude of variations
V.5.2. δ13C depletion
V.5.3. Phase differences
V.5.4. Existence of a seesaw pattern
V.6. Sequence of events during TII
V.7. Clathrated ice and TII results
V.8. Conclusion
CHAPTER VI. CONCLUSIONS
VI.1. Technical conclusions
VI.2. Summary of the new key dataset
VI.3. Conclusions on the carbon cycle / climate interactions
VI.4. Perspectives
REFERENCES
APPENDICES
