Air-Sea Gas Exchange
An introduction essay
Andreas Andersson
Department of Earth Sciences, Meteorology
Villavägen 16
752 36
Uppsala, Sweden
1
Contents
1 Introduction ........................................................................................................................................ 3
2 Spectra................................................................................................................................................. 4
3 Co-spectra ........................................................................................................................................... 7
4 Flux measure and estimation techniques .......................................................................................... 8
4.1 The Eddy Covariance method ......................................................................................................... 8
4.2 Footprint and internal boundary layers........................................................................................ 10
4.3 The Bulk method........................................................................................................................... 11
5 The carbon cycle and the importance of oxygen ............................................................................. 12
5.1 Oxygen .......................................................................................................................................... 13
6 Gas exchange at the Air-Sea interface ............................................................................................. 15
6.1 Transfer velocity ........................................................................................................................... 16
6.2 Solubility ....................................................................................................................................... 17
6.3 Processes influences transfer velocity .......................................................................................... 19
6.3.1 Sea spray and bubbles ........................................................................................................... 19
6.3.2 Water-side convection ........................................................................................................... 20
6.3.3 Precipitation ........................................................................................................................... 21
6.4 Processes of importance for ∆pCO2.............................................................................................. 22
6.4.1 Upwelling .............................................................................................................................. 22
7 The COARE-Algorithm ....................................................................................................................... 23
7.1 Surface renewal theory ................................................................................................................ 23
7.2 COARE (2.0, 3.0) algorithm ........................................................................................................... 23
8 Wintertime Air-Sea fluxes of CO2 ..................................................................................................... 25
8.1 Air-Ice -Sea interaction ................................................................................................................. 25
9 The Baltic Sea and the site of Östergarnsholm ................................................................................ 28
9.1 Östergarnsholm ............................................................................................................................ 29
10 MicroTX3 oxygen sensor................................................................................................................. 30
11 Future work ..................................................................................................................................... 32
2
1. Introduction
Over 70% of the Earth surface is covered by ocean and it contain over 50 times more
inorganic carbon than the atmosphere. In the history of Earth the pCO2 level in atmosphere
has been varying a lot. At the start of industrialization of CO2 the global atmospheric CO2
level were steady at 280μatm and the air-sea net flux of CO2 was approximately zero. Today
the fluxes of CO2 across the air-sea interface are enormous and averages about 1.5- 2.0
GtonC/yr into the ocean (Takahashi et .al, 2002, 2008) and the global CO2 concentration
averages about 390 μatm and is steadily increasing. It has been known for several years that
this large flux of CO2 causes acidification of the world’s ocean, but in the last 20 years more
effort has also been put into the effects from the enhanced emissions due to human activity, in
terms of global warming. For instance the climate panel, IPCC (Intergovernmental Panel on
Climate Change) was founded in 1988 to summarize the current knowledge and ongoing
research in assessment reports. According to the latest report IPCC ( 2007), the predicted
doubling of the CO2 concentration during the next hundred years, will result in an increased
global mean temperature by 2°C - 4.5°C. In (IPCC, 2007) it was stated that the global net
flux of CO2 is directed into the sea, but the regional and seasonal distribution are still
uncertain for some areas. To make a more detailed description, determine areas of sinks and
sources of CO2, better knowledge about global carbon cycle and processes involved are
needed. In order to map what areas of the World´s Ocean are a sinks or sources of CO2, and
thereby the impact of the anthropogenic enhanced emissions of CO2 on the ecological system,
a good parameterization of the air-sea exchange of CO2 is needed.
The magnitude and direction of the flux of CO2 across the air-sea interface is governed by the
difference in partial pressure (
) between the water and atmosphere and by the transfer
velocity, which describes the efficiency in the transport process. The most commonly used
equations to describe the fluxes of gases across the air-sea interface uses the transfer velocity
as a function of wind speed only (Liss and Merlivat, 1986; Wanninkhof, 1992; Wanninkhof
and McGillis, 1999; Weiss et al., 2007) in linear, quadratic or cubic form. Since these
expressions for the transfer velocity is a function of wind speed only they become sensitive to
changes in wind variability. Takahashi et al. (2002) found after applying the cubic wind
dependent new wind dependent function for transfer velocity (Wanninkhof and McGillis,
1999) on a global model, a 70% increase in the oceanic uptake of CO2, compared to what was
found for the quadric wind dependent function. However there are other physical processes
affecting the transfer efficiency across the air-sea interface such as: biological production,
upwelling, spray and bubbles (Woolf, 1993, 1997), water-side convection (Rutgersson and
Smedman, 2010) micro scale wave breaking (Zappa et.al, 2001) and the cool-skin effect
(Robertson and Watson, 1992). In coastal and shallow regions were upwelling and biological
production have a major affect on the air-sea exchange of CO2, these global models fails,
periods of spring bloom or with low wind speeds are often left out. This because of missing
detailed knowledge about important parameters controlling the air-sea exchange of CO2.
In the biological production oxygen plays an important role as one of the end products in the
photosynthesis and since O2 is less soluble in water than CO2, oxygen serve as a more
3
effective tracer gas for biologic activity than CO2 (Keeling,1993). In a study made in the
Baltic Sea Rutgersson and Smedman (2010), showed that during periods with a well defined
mixed layer, a significantly enhanced transfer velocity was found due to surface cooling
resulting in water-side convection. A significant seasonal cycle for pCO2 in water, with peak
values above 800 µatm in wintertime and as low as 150 µatm during summer were also seen.
It has also been shown that the effect from precipitation alone, in areas characterized by low
wind speeds (The Western Pacific), are of the scale of changing the annual flux of CO2 both
by magnitude and direction, from being a source of CO2 to being a sink of CO2. During
rainfall events the surface
has seen to be lowered by as much as 30µatm, caused by
chemical dilution from rain. After including the modeled effect from rain on the air-sea CO2
exchange, the Western Equatorial Pacific becomes as sink of -0.078 mol CO2 m-2 yr-1 instead
of a ocean source of +0.019 mol CO2 m-2 yr-1 (Turk et al., 2010).
In high latitudes the largest sinks in the global CO2 global During wintertime in high latitudes
large temperature difference between air and water exists, causing deep convection and severe
weather. For the air-sea exchange of gases the ice extent becomes an important feature.
Traditionally the ice has been considered as an impermeable for gas exchange, but recent
studies shows that there exist a direct exchange of CO2 between sea ice and atmosphere
(Papadmitriou et al, 2004; Delille et al., 2007; Diekmann et al., 2008 ; Loose and Schlosser,
2011; Papakyriaiou and Miller, 2011). Springtime measurements over ice (Papakyriaiou and
Miller, 2011) showed CO2 fluxes comparable to those found over open water, with a
maximum hourly efflux of CO2 of 1.0 µmol m-2 s-1 and -3,0 µmol m-2 s-1 into the ice. On
average an efflux of 0.36 µmol m-2 s-1 was found.
2. Spectra
Energy spectra
relates the size of turbulent eddies ( to the energy of the specified
quantity like temperature, humidity, CO2 and other scalars. According to Monin-Obukhov
similarity theory three different regimes in the energy spectra can be identified (figure 1). The
first is the energy containing range or production regime, where the energy is transformed into
turbulence from the mean flow. In the next regime, inertial subrange (0.01-5Hz), there is no
production of turbulence and the flow is considered isotropic, which is not the case for large
eddies in the energy containing range. In the inertial subrange large eddies cascades into
smaller eddies, with negligible viscosity forces, all kinetic energy is conserved. In the last
regime, the viscous subrange, small eddies breaks down into thermal energy by viscous
forces.
Turbulent eddies are often described as circular motions, for which the velocity u is defined as
function of the radius r of the motion trajectory according to:
(1)
It has been shown (Sutton, 1953) that the friction force can be expressed like:
4
(2)
Where is the wind velocity and is the kinematic viscosity. The energy dissipation per time unit
is equal to the friction force multiplied by the velocity By assuming that the specified velocity is
equal for all sizes of eddies, the dissipation becomes
(3)
The dissipation is inversely proportional to the third power of the radius of the eddies
,
thus small eddies becomes much more efficient than larger eddies of breaking down turbulent
energy. Since small eddies by definition have a short length scale their characteristic time
scale also becomes small, causing them to respond quickly to a change in the mean flow
(Högström and Smedman, 1989), thus in equilibrium with the mean flow. According to the
theory of Kolmogorov the shape of the energy spectra than only depends on the velocity by
which turbulent energy is fed from lower frequencies towards higher frequencies. Since no
energy is released during the transformation of turbulent energy within the inertial subrange,
the energy spectra
within this regime can be expressed as a function of the dissipation
rate
and the wave number in x-direction
:
(4)
a and b are dimensionless constants, which can be solved by dimensional analyze, then gives
the formula
(5)
Figur 1. Energy spectra, including the different regimes (Högström and Smedman, 1989).
5
Where
denotes the universal Kolmogoroff constant for
equal to 0.52 (Högström 1996). By
applying the Taylor hypothesis the wave number can be expressed in terms of mean wind
velocity and frequency (n), makes the equation (5) to look like:
(6)
In order to compare measurements made at different sites, where time height and atmospheric
stability are allowed to vary, a normalizing of spectra is needed. One commonly used way to
normalize spectra is by using the similarity theory of Monin-Obukhov and following the steps
presented in Sahlée (2007), where the Obukhov length, L
dimensional frequency
.
is the friction velocity,
, and the non-
the von Karmans constant equal
to 0.4 (Högström, 1996),
the temperature and
is the vertical virtual temperature flux.
Then eq. (6) is normalized with
and the energy spectra becomes a function of the
non-dimensional stability parameter z/L.
(7)
where
is the dimensionless dissipation of turbulent energy, thus a function of z/L .
To make the spectra independent of stability eq. (7) is divided by
, and takes the form
(8)
Kaimal (1973) showed that by normalizing the spectra with the controlling parameters similar
as above, all spectra and co-spectra coincide and follow a set of universal curves. Corrsin
(1951) derived an equation for the temperature spectra, , with the assumption that an inertial
subrange also exists for temperature.
(9)
In similar way spectra for other scalars such as humidity and carbon dioxide can be derived
(10)
Where
and
are dissipation rates for temperature and other scalar variances respectively,
and
are dimensionless constants, showed to be the same for all scalars,
=0.8 by
Hill (1989). However Norman et al. (2011) found = 0.68 by using direct flux measurement
at site Östergarnsholm. As for eq. 5 Taylor hypothesis can be applied and normalized with the
6
scaling variables, which for
becomes
and for
:
. By removing the
stability dependence as for equation 8, equations 9-10 takes the form:
(11)
(12)
3. Co-spectra
The covariance is the average of the product of the two fluctuating variables. When
determining the vertical turbulent fluxes in meteorology by Eddy Covariance technique,
correlation between the fluctuation part of the parameter ( ) and the vertical wind component
(wˈ) is used. The co-spectra is defined as the real part of the cross-spectra. To get the cospectra for
one first needs to define the cross-covariance between w and c, in order to get
the cross-spectrum.
(13)
Where
is the cross-covariance and the separation distance between were the parameter
and the vertical wind is measured. The cross-spectrum can then be obtained by taking the
Fourier transform of the cross-covariance
(14)
Where the real part of the cross-spectrum
becomes the co-spectrum, and the
imaginary part is named the quadrature spectrum and represents the part of the cross spectra
that is out of phase. The result when computing the integral of the quadrature spectrum will
then be identically to zero, thus having no contribution to the flux (Lenschow, 1995). The flux
can be determined by computing the integral of the co-spectrum.
(15)
If the covariance
takes positive values the two scalars are positive correlated and if the
covariance become negative they are negatively correlated. In order to get information about
the fluxes the co-spectrum has to be weighted and normalized, usually y-axis with product of
the fluctuating scalars
and the x-axis with z/U. The co-spectrum gives information about
which sizes of eddies (k) is contributing to the flux. In a semi-log presentation the area
beneath the curve between two frequencies directly corresponds to amount of the flux that is
transported with that frequency, while a linear plot positive y-values corresponds to positive
(upward) fluxes and negative y-values to downward fluxes.
7
4 Flux measure and estimation techniques
There exist different measurement techniques in order to determine the turbulent fluxes of
impulse, temperature and other scalars. Common for most of the measuring techniques is that
they are based on different assumptions and simplifications.
4.1 The Eddy Covariance method
The Eddy-covariance method (also called eddy-correlation method) is a direct method where
no applications of empirical constants are needed in order to describe the turbulent fluxes
(Foken et al., 1995, Kaimal and Finnigan 1994). In this method measurements of the wind
components are done with a high frequency instrument, usually a sonic anemometer. Then
one can achieve the turbulent fluxes of momentum and with an additionally instrument also
fluxes of gases such as CO2, CH4 and O2. To get the true fluxes the averaging period has to be
long enough so that the energy contributions from all frequencies are correctly represented. If
the averaging period is to short the low frequency contribution to the fluxes are missed or not
correct represented. The shortest averaging period is dependent on height above ground and
stability. For heights of 2-5 m a averaging period of 10-20 min would be required for daytime
unstable stratification. The size of the turbulent eddies are dependent of the height above
ground, were the size of the eddies increases with height. Therefore the ideal measuring path
length and the separation length between the sonic anemometer and the additionally device
also depends on height above surface. For a 12cm device the measurement height should not
be below 2m above surface (Foken, 1995).
To apply the eddy covariance method, the assumption of negligible mean vertical wind has to
be full filled. This is done by tilting the horizontal plane into the mean wind direction so that
the vertical component of the mean wind vector becomes. This rotation is done in two steps
by using the measured wind components (subscript ). The first step is the rotation of the
coordinate system around the z-axis into the mean wind.
cos
(16)
(17)
(18)
Where u and v denotes the horizontal wind components in x- and y-direction, respectively, w
is the vertical wind and is defined as:
(19)
8
The second rotation is to rotate the coordinate system around the new y-axis until the mean
vertical wind disappears (Kaimal and Finnigan 1994).
cos
(20)
(21)
sin
Where
(22)
is defined as:
(23)
The ideal gas law states that for constant pressure, a rising temperature corresponds to a
expanding volume. Since the instrument is measuring the concentration of a certain
substance, fluctuations of temperature and water vapour causes the detection volume to
contract and expand, which changes the measured density of the substance, which are not true
surface fluctuations. In order to get the true fluxes out from those due to changes in
temperature and moist, a separation method is needed. The most common method (Webb et.
al., 1980) is presented below
(24)
Where the second term denotes the compensation for the latent heat flux and the third term
compensate for the sensible heat fluxes. In recent years other methods for removal of the
contribution from fluctuations of temperature and moist to the true flux has been developed,
such as by Detto and Katul (2007) and the Direct Conversion-method (Sahlée et al., 2008). I
the DC-method method, the density signal is directly converted into mixing ratios. Below the
theory and the step of the DC-method is presented.
(25)
where is the measured air pressure and
calculated by using:
the partial pressure of water vapour, which can be
(26)
Here
is the molar densities of water vapour (mol m-3 ),
the dry gas constant 287.06 (J
-1 -1
kg k ) ,
is the molar mass, the subscript and denotes water vapour and dry air
respectively, and
is the ratio
.
Then the mass density
(kg m-3) and the molar density of dry air
calculated by using the Gas law:
9
(mol m-3) can be
(27)
(28)
Finally the mixing ratios of CO2 and water vapour can be calculated:
(29)
(30)
Following these steps the true fluxes of water vapour and CO2 can be achieved by Reynolds
averaging and then use eq. (13-15).
4.2 Footprint and internal boundary layers
Measurements made at a certain site time and at a given height does not represent the
properties of the fluxes at the place of the instrument. Instead the properties of the
measurement corresponds to the conditions of the underlying area upwind the site. This area
is called the footprint and the size of it depends on the wind vector, atmospheric stratification,
measurement height and surface roughness.
To use the eddy covariance method there are some limiting conditions, which need to be
fulfilled such as horizontally homogeneous surface and steady state condition. Therefore one
should ensure that the footprint area is a uniform area for all stability conditions. If the
footprint constitutes an area larger than the uniform area, which were supposed to be
represented in the measurements, different turbulent structures from the surroundings will be
present in terms of internal boundary layers (fig. 2), disturbing the measurements. For
measurements close to the surface layer it´s then essential to identify internal layers and their
structure.
Fig 2. (Stull, 1988). Generation of internal boundary layer above an inhomogeneous surface
Obstacles such as buildings and trees have big influence on meteorological measurements and
the disturbance is not only concentrated to the downwind side but also on the windward side.
10
Therefore information from wind tunnel simulations constitutes a good supplement, even
though wind tunnel simulations are limited to neutral stratifications and atmospheric
turbulence only can be provisionally created. The best thing to do is to choose a site were the
influence of obstacles can be excluded as much as possible.
4.3 The Bulk method
The Bulk method is one of the profile methods, were the flux-gradient similarity is used. In
the Bulk-method a uniform linear gradient is assumed, thus the most simple profile method
for energy exchange. The bulk approach is sensitive to internal boundary layers, uniform fetch
are required, which makes it almost only applicable over water.
The turbulent fluxes of momentum, sensible heat and latent heat are related to the gradients
via the bulk coefficients: drag coefficient Cd, Stanton number Ch, and the Dalton number Ce.
(31)
(32)
(33)
Where is the potential temperature, q is the water vapor mixing ratio (usually at 10m height)
and is one of the horizontal wind components components relative to the fixed earth and S
the average of the wind speed relative to the sea surface at the reference height .
is the
sea surface interface temperature;
the surface current and
the interfacial value of the
water vapor mixing ratio, computed from the saturation mixing ratio for pure water at the sea
Surface Temperature (SST) (Fairall et al.,1996). By using the Monin-Obukhov similarity
theory and express the scalar and wind component as non-dimensional profile functions, the
exchange coefficients become:
(34)
(35)
Where ,
,
, are the roughness length for momentum, temperature and humidity
respectively, and k is the von Karmán constant equal to 0.4 ( Högström et al., 1996) .
, ,
are the integrated form of the non-dimensional gradients, where the non-dimensional
gradients have the form
(36)
(37)
11
(38)
where,
is the friction velocity,
and ,
denotes the scaling
parameters for temperature and humidity, and z/L is the Monin-Obukhov stability parameter,
where L is the Obukhov length defined as:
, where
is the potential virtual
temperature.
In general the flux coefficients are stability normalized by using a reference neutral
stratification for all coefficients, and by setting the reference height z=10m. In order to
compare the fluxes of sensible and latent heat to measurement preformed at a different site
and another atmospheric stability, the stratification influence is removed, then CE and CH are
transformed into their neutral counterparts CEN and CHN.
(39)
(40)
5 The carbon cycle and the importance of oxygen
The carbon cycle is the biochemical cycle which describes the exchange of carbon inside the
biosphere between geosphere, pedosphere, hydrosphere, atmosphere, including the fluxes of
carbon dioxide between ocean and atmosphere. The carbon cycle is perhaps the most
important cycle found in nature, since it contains the spine of all organic substances through
chemical reactions with other elements. Usually the carbon cycle is divided into two different
cycles, the geological carbon cycle, with a time scale of millions of years and the
biological/physical carbon cycle (fig.3), which operates over shorter time scales (days to
thousands of years) (NASA, 2011). In figure 3 the contribution in Gton carbon is displayed
for the reservoirs (white), while natural fluxes (yellow) and the fluxes from human emissions
(red) between the reservoirs are in Gton/year. There it’s stated that the ocean acts as a sink of
carbon dioxide by 2Gton carbon/year due to human emissions, while at the time of the start of
industrialization the air-sea net flux were equal to zero.
12
Figure 3. The Carbon Cycle (Riebeek,2011), Illustrating storage and fluxes of CO2 gigatons of
Carbon and gigatons of Carbon/year. Yellow numbers are natural fluxes, red are anthropogenic
fluxes and white numbers denotes stored carbon.
5.1 Oxygen
Oxygen plays an important role in the global carbon cycle, where oxygen by chemical
reactions can be either produced or consumed. The most important chemical reaction is the
photosynthesis, where carbon dioxide reacts with water under the absorption of sunlight to
form organic matter as carbohydrates (sugar) and oxygen. The reversed chemical reaction
where carbon dioxide and water are produced during breakdown (oxidation) of organic matter
is the respiration. In one year the amount of carbon taken up by photosynthesis from the
atmosphere and released back by respiration is 1000 times greater than what moves in the
geological cycle. A simplified described scheme of photosynthesis and respiration in
terrestrial biota is showed as
(41)
The ratio of O2 and CO2 averages about 1.05 (Keeling, 1988), thus photosynthesis exceeds
respiration, resulting in that organic matter slowly builds up and forming oil and coal deposits. In the
marine biota photosynthesis respiration reaction looks as
(42)
13
represents the approximate composition of marine organic matter
(Redfield et al. 1963), thus the actual reaction depends on the composition of the organic
matter. Then there also exists the anthropogenic burning of fossil fuels.
Where
(43)
These are the dominant reactions causing sinks and sources of atmospheric O2 and CO2. The
geochemistry of CO2 differs from O2 and there are additional chemical reactions in water, which
affects pCO2 in water and by that the air-sea exchange process and the concentration of atmospheric
CO2. The time scale for the upper ocean to equilibrate with the atmosphere with respect to CO 2 is in
the order of a year or more, while in the case for O2 it´s only a matter of a few weeks (Broecker and
Peng, 1974, 1982), (Keeling, 1993). For example CO2 reacts with water according to eq. (44) and
form carbonic acid, which then reacts stepwise into carbonate ion.
Conversion to carbonic acid:
CO2 (dissolved) + H2O ⇌ H2CO3
(44)
First ionization:
H2CO3 ⇌ H+ + HCO3− (bicarbonate ion)
(45)
Second ionization:
HCO3− ⇌ H+ + CO3− (carbonate ion)
(46)
For every mole fraction of CO2 two moles of the hydrogen ion (
) are produced, causes the ph-level
in the ocean to drop. The different behavior of CO2 and O2 gives that the ocean serve as a major
reservoir of CO2 but only a minor reservoir for O2.
In subpolar oceans were the commonly used definition of the depth of the mixed layer (ΔT=0.5°, and
density Δ =o.125kgm-3) proposed by Monterey and Levitus (1997), are not suitable in winter time,
leading to an overestimation of the mixed layer depth (Kihm and Körtzinger, 2010), instead oxygen is
used as a tracer to determine the depth of the mixed layer (
) (Körtzinger et al.,
2008). During wintertime, in the central Labrador Sea, the deepest convection occur forming a mixed
layer down to 1350m according to the definition above. The fast deepening of the mixed layer results
in an unfulfilled equilibrium state of oxygen between atmosphere and ocean typically 5-7%
undersaturation (Kihm and Körtzinger, 2010). The spring bloom then causes a production of oxygen
and depletion of the nutrition´s solved in winter time, leading to a strong supersaturation (~15%).
Later on the air-sea exchange gradually reduces the oxygen concentration within the sea. The
generally low wind speeds and heating of surface water in summer time generates a shallow mixed
layer. Underneath the mixed layer, a layer of subsurface water is found where a major part of the
summer production in the overlying waters is reminilized via respiration. This forms a large oxygen
14
gradient between surface waters and the respiration formed oxygen deficit layer. This deficit layer is
of major importance for the strong oxygen flux into ocean later on, during the deep convection period
(Kihm and Körtzinger, 2010). The different oceanic behavior of CO2 and O2 makes that variations in
O2 gives information about the carbon cycle that is not contained in variations of CO 2, leading to one
important application of O2: Measurements of O2 gives a more precise rate of the biological new
production in the ocean than measurements of CO2.
6 Gas exchange at the Air-Sea interface
In all aqoues environments such as sea, lakes and rivers gas exchanges across the air-water
interface occur. Most of the parameterizations relies on the two layer film model (Liss and Merlivat,
1974), where the exchange zone is separated into two layers one in air and the other one in water,
closest to the surface. These two layers are assumed to be in equilibrium with each other and the
transport across the air-sea interface is determined by molecular diffusion. The exchange of a scalar
across that surface is then depending on the fugacity and the efficiency of the transfer process, where
turbulence are of most importance (MacIntyre et al., 1995). The fugacity is the tendency for the gas to
escape from its present state, or so to say the effective pressure. The fugacity is equal to the pressure of
an ideal gas which has the same chemical potential as the real gas, for example N2 at 0°C and pressure
of 100 atm. has the fugacity of 97.03 atm. This means that the chemical potential of real N2 at the
pressure of 100 atm. has the same value as ideal N2 having the pressure 97.03 atm. (Smith, 2004;
Atkins and Atkins, 2008). In accurate chemical equilibrium calculations the fugacity is considered
instead of the mechanical pressure, which are related to each other by fugacity constant .
(47)
However for
under normal conditions in sea,
differs from
by less than 0.5%
(Skjelvan et al., 1999), thus
usually are considered instead of the fugacity, and the flux across
the air-sea interface can be expressed as:
(48)
where k is the transfer velocity, K0 the solubility and
=
. The transfer
velocity is in most cases considered to depend on the wind speed (U10) and the Schmidt number (Sc).
However there are several other parameters affecting the transfer velocity and
in water,
that from time to time are capable of governing exchange of
across the air-sea interface,
and thereby change the flux both by magnitude and direction, such as: water-side convection,
precipitation, bubbles caused by sea spray, upwelling, biological production, cool skin effect
and atmospheric stability (fig. 4). These features either affect the transfer velocity or the
actual
. In global models many of these parameters are often left out, because of lack of
knowledge, thus no satisfying parameterization can be presented. Parameters as bubbles
(Woolf, 1993) and water-side convection (Rutgersson and Smedman, 2010) has been
described in terms of a contribution to the transfer velocity.
15
Figure 4. Modified from (Wanninkhof et al., 2009). Simplified schematic picture, displaying factors
affecting the air-sea exchange of CO2
6.1 Transfer velocity
In the literature, linear, quadratic and cubic empirical relations have been suggested, in how to relate
transfer velocity to wind speed. A quadratic relation suggested by Wanninkhof (1992) is frequently
used:
(49)
where U10 is the horizontal wind speed at 10m and Sc the Schmitt number, in this case normalized
with Sc=660 (CO2 at 20°C in seawater ).
Turbulent eddies exists in a broad range of sizes, where the largest scales with the height of
the boundary layer and the smallest eddies can be down to the Kolmogorov microscale, which
in atmosphere is about
m (Kaimal and Finnigan, 1994). These turbulent eddies are
causing a mixing both in air and water, resulting in a diffusion of mass all the way through the
mixed layer of the water column. In the two layers (fig 5) just above and below the air-water
interface, viscous forces dominate and thereby suppresses turbulent mixing, causing the
molecular diffusivity to dominate. At this stage the Schmidt number is a useful term as it
relates the diffusive sublayer to the kinematic viscosity of the water and molecular diffusivity.
The Schmidt number is the ratio of the kinematic viscosity of seawater (m2/s) to the mass diffusion
coefficient of the considered gas
.
16
(50)
Where the mass diffusion is described as a function of temperature:
(51)
is the maximum diffusion coefficient, EA the activation energy, T is temperature (K) and R
is the gas constant. The dependence of the Schmidt number on the gas transfer velocity opens
up the opportunity to relate gas transfer velocity of one gas to another. If the transfer velocity
of gas 1 is known, the transfer velocity of gas 2 can be computed by using the Schmidt
number for the respective gas according to:
(52)
Usually the gas transfer velocities are normalized to Sc=660, which is the Schmidt of CO2 at
20°C in seawater, or Sc=600 (CO2 at 20°C in freshwater) (Bade, 2009)
Another commonly used formula for air-sea transfer of CO2 flux is the bulk aerodynamic
formula:
Δ
(53)
DCO2 denotes the bulk transfer coefficient, related to the transfer velocity by: DCO2=k660/U10, where
U10 is the wind speed at 10 meters height.
6.2 Solubility
The efficiency of the transfer process in the air-sea system of a gas (n) is as previously stated
a function of partial pressure (pn ). The solubility in atmosphere and water for different gases
differs by several orders of magnitude. The boundary regions in the exchange layer between
air and water are separated into different layer. Closest to the interface is the aqoues diffusive
sublayer and the diffusive layer of air. The upper most part of the aqueous diffusive sublayer
is in equilibrium with air. The profile through the boundary region differs between gases
depending on solubility. For soluble gases such as H2O, SO3 and NH4 the exchange ratelimiting step is across the boundary layer in the air, while for slightly soluble gases such as
CH4, DMS, CO2 and O2 the retardation is in the aqueous boundary layer (Matson and Harris,
1995).
17
Figure 5 ( Waninkhof et al.,2009). The boundary region in the exchange layer, and vertical
profiles through the exchange layer for gases with different solubility. A slightly soluble gas
(blue) where the exchange rate limiting part is in the aquatic boundary layer, while for a
soluble gas (red), the limitation is found in the air boundary region.
The osmotic pressure of a atmospheric gas (n) is determined by, concentration, temperature,
according to eq. 54.
(54)
where Rn is the gas constant for the specified gas n. The solubility ( ) is defined as the
greatest possible amount of the gas (n) which can be dissolved in a unit volume of liquid
(Krauss and Businger,1994), and can be determined according to:
(55)
P0(n) is a reference pressure in the ambient gas phase, mostly used 1013,25hPa. Pn denotes the
saturation vapour pressure of the gas (n) which would be in equilibrium with its condensed
phase. For oxygen pn equals 36,000 standard atmospheres at 0°C. Saturation pressure always
decreases with falling temperatures, causing the solubility of the gas (n) to be a function of
temperature depth and latitude (Krauss and Bussinger, 1994) . In addition the solubility of the
gas(n) is also dependent on the nature of the solvent which is included in the proportionality
constant in eq. 55. Sea salt and other heavy substances in the solvent, decreases the solubility
18
of the gas known as the Setschenow effect. Considering a given solution where the gas (n) is
included, adding a heavy salt molecule, results in a lowering of the mole fraction of the gas
(n) in the solution, and thereby decreases the pressure ratio in eq. 55, causes the solubility to
decrease. Looking at a specific gas and going from fresh water to seawater the solubility of
the gas decreases by approximately 20%.
However in the atmosphere the partial pressure of the gas Pn is necessarily smaller than total
pressure p0. Therefore will the standard value of the saturation concentration ( ) of the gas
(n) in surface sea water be described by
n
(56)
Although in a submerged air bubble, the partial pressure pn is larger than in surface air, causes
the saturation concentration to increase with water depth (Krauss and Bussinger, 1994).
6.3 Processes influences transfer velocity
6.3.1 Sea spray and bubbles
Sea spray occurs when a wave is mechanically disrupted due to strong winds or by bursting of
bubbles at the sea surface. When a bubble collapses its surface free energy is converted to
kinetic energy that rises like a vertical jet from the centre of the cavity. Eventually this
vertical jet breaks up into 1-10 drops (Woolf, 1993)
The structure of the air-sea interface is mainly determined by wind speed. Bubbles are formed
when air has been mechanically entrapped in water, mostly due to breaking waves seen as a
whitecap coverage. During breaking wave’s situation a layer of captured bubbles will form
below the surface. This layer is called the bubble cloud, within this bubble cloud the
encaptured air interacts with the sea water. Then depending on the relation of partial pressures
(
) and chemical reactions, gases and materials could then either be dissolved into
the sea, or picked up by the air and released into the atmosphere.
Releasement of aerosols into the atmosphere as a consequence of whitecap cover and wave
braking is known to have a positive effect for cloud formation. It has also been suggested by
Woolf (1993) to have an important effect on the air-sea fluxes (F) of greenhouse gases can be
described by.
(57)
19
Where Kb is the bubble contribution to the transfer velocity, K0 the transfer velocity for direct
exchange S the solubility, p the partial pressure of the gas, C the concentration and Δ denotes
the equilibrium supersaturation.
For the total air-sea fluxes bubbles seem to have a much smaller effect for more soluble gases
such as CO2, for poorly soluble gases as methane. However the bubble transfer velocity
relationship on the Schmitt number and solubility of the gas depends on the bubble size
distribution as well as their vertical mixing. In addition the bubbles hydro dynamically clean
or dirty state is of importance. A dirty bubble refers to a bubble, which has a surface
containing active material, compared to a clean bubble, characterized by a completely free
surface of active material. This surface-active material has a negative effect on the rise time of
the bubble as well as the bubble transfer velocity (Kb) (Woolf, 1993). As stated earlier
bubbles are generated from breaking waves, thus bubbles are a function of wind speed.
Therefore it’s common to use the whitecap coverage as an appropriate index for the strength
of the bubble contribution to the total transfer velocity.
6.3.2. Water-side convection
A mixed layer depth is usually defined as the depth, at which the density is about the same as
at the water surface. Inside the mixed layer water column the properties of density,
temperature and salinity are more uniform, due to the mixing. The depth of the mixed layer
varies during the year, for most places. The mixing arises from waves and turbulence
generated at the sea surface, in most cases caused by the by the wind stress eq. (33).
Stress-induced mixing inside the water-side mixed layer is mainly constrained to intermediate
to high wind speeds. For low to intermediated wind speed convection is an important feature,
causing mixing inside the water column and thereby enhances the vertical transport of
dissolved gases like CO2, resulting in an enhanced turbulence at the air-sea interface (Eugster
et al., 2003). Convective mixing as a result from surface cooling can occur due to a cold air
mass advection or as a result of a diurnal cycle. Macintyre et al. (2003) and Jeffery et al.
(2007) defined an expression for the strength of the water side convection:
(58)
Where is the convective velocity,
the depth of the mixed layer and
buoyancy flux defined by Jeffrey et al. (2007) as
the water-side
(59)
Where g is the acceleration by gravity, the thermal expansion coefficient,
the net
surface heat flux,
the specific heat capacity of water,
density of water, is the saline
expansion coefficient,
the latent heat flux and is the latent heat of evaporization.
20
By separating a time serie into two periods: period 1 spring (well defined mixed layer > 20m)
period 2 summer (shallow mixed layer), Rutgersson and Smedman (2010) showed a
significant enhancement of the transfer velocity during period 1 and unstable atmospheric
stratification. Where the magnitude of the enhancement increased as the surface cooling
increased, and thereby the water-side convection, thus a modification of eq. (49) was
suggested, where the impact of the water-side convection
is included.
(60)
Where
is defined as a function of the convective velocity
(61)
6.3.3. Precipitation
Recent studies have shown that rain may have a significant effect on air-sea exchange of CO2
and other gases (Takagaki, 2007 and Turk et.al, 2010). There are several known processes
taking place in the water surface layer during rainfall, even though the extent of each process
still is uncertain. Surface chemical dilution caused by rain reduces the near surface salinity,
total alkalinity and dissolved inorganic carbon, and thereby decreases the pCO2 in the sea
surface water (Dickson et.al, 2007, Turk et.al, 2010). High wind speeds will result in an
increased vertical mixing, acting to homogenize these gradients. In situations with rain and
low wind speeds this surface dilution effect can be maintained for a significant time.
The second effect is the rain-induced surface turbulence (RIST). This turbulence causes a
shallow vertical mixing in the water column, thus enhances the transfer velocity [Ho et al.,
1997, 2000, 2004; Takagaki and Komori, 2007; Zappa et al., 2007, 2009]. A third effect is the
wet deposition of dissolved inorganic carbon (DIC) (Komori et.al., 2007), dissolved organic
carbon (DOC) (Willey et al., 2000) and the composition of the rainwater drops, making it
hard to study all these effects together. Laboratory experiments (Zappa et al., 2009) has
showed that the rain induced surface turbulence caused a substantial increase in the transfer
velocity (k). Also a shallow stable stratified layer formed due to the addition of cold fresh
water, thus trapped the RIST very near the surface 10-20cm. In Henocq et al, (2010) the rain
induced effect on the vertical salinity gradient was estimated in the top 10m, using mooring
measurements from the Tropical Atmosphere Ocean project. There it was shown that
measurements at 5m depth often do not capture the dilution due to rain that could be detected
at 1m depth.
According to recent studies (Turk et al., 2010) the sum of the rain induced effects changes the
oceans role in the global carbon cycle for some areas. The Western Pacific, characterized by
low wind speeds during rainfall and a pCO2 of 10-30 µatm above the atmospheric , after
21
including the effect from rain becomes as sink of -0.078 mol CO2 m-2 yr-1 instead of a ocean
source of +0.019 mol CO2 m-2 yr-1.
6.4 Processes of importance for
Setting the transfer velocity equal to one, the gas exchange between two reservoirs becomes
controlled by the difference in partial pressure of the specific gas. For the state of the air-sea
the governing processes, capable of changing the flux by magnitude and direction are
mainly found within the water, and vary both in space and time. Biological activity (section 5)
is considered to be the strongest parameter, lowering
in spring and summertime.
On shorter time scales other processes also becomes important such as upwelling, dilution of
surface water by rain, changing of sea surface temperature and ice formation.
6.4.1 Upwelling
Upwelling is a wind generated phenomena, and during upwelling conditions, deep water is
vertically mixed with nutrient surface water. Upwelling can be seen both in the synoptical and
local scale, but it’s usually divided into three main types: coastal, equatorial and seasonal
upwelling (NOAA, 2011). The intensity of the upwelling, are besides wind strength also
dependent on the vertical structure of temperature and salinity in the water, altitude and the
structure of the sea bottom.
The photosynthesis act to bind CO2 into organic matter, some of it sink down to the
deepwater, forming sediments, and some of it stays in the surface layer where it can be
released back to the atmosphere. During decomposition of organic and inorganic carbon in the
surface layer as well as in the deepwater, CO2 is released. In case of an upwelling situation,
high levels of CO2 from deep water can be brought to surface and released to the atmosphere.
This water-side turn around caused by upwelling, triggers a convective mixing, enhancing the
vertical mixing inside the water column and the exchange of gases across the air-sea interface.
In situations with Strong winds, upwelling, or surface cooling as a result of cold air advection,
a vertical mixing inside the water column also occurs, and CO2 saturated surface water is
replaced by unsaturated water from deeper layers. This cycle can be enhanced in case of an
existing deep water circulation, as a result from currents and winds.
In terms of air-sea fluxes of CO2, a region of upwelling of cold water generally enhances the
downward fluxes of CO2 into the ocean (Hales et.al, 2005). However in summertime when the
normal CO2 flux is directed downward, a situation with upwelling brings up colder and more
CO2 containing water to the surface, capable of changing the flux both by magnitude and
direction.
22
7. The COARE-Algorithm
In 1996 the first version of the Coupled Ocean-Atmosphere Response Experiment (COARE
2.0) was published (Fairall et al., 1996). Since then, two updated versions have been presented
COARE 2.5 and COARE 3.0 (Fairall et al., 2003). Today the COARE algorthim is one of the
most frequently used algorithms for calculation of air-sea fluxes in the scientific community.
The COARE algorithm is based on the surface renewal theory and uses a bulk algorithm (see
section 2.3) to calculate the fluxes, the latest version is considered as state of the art. A briefly
description of the theory and the involving steps in the COARE algorithm is presented in
section 7.2.
7.1 Surface renewal theory
The basis of the surface renewal theory (Brutsaert, 1975) is that, the exchange at the air-sea
interface is governed by the diffusive transport by small turbulent eddies across a thin
interfacial sublayer, approximately 1mm thick. This thin layer is composed by the sub layers
in air and water. Within the interfacial layer the fluxes of sensible and latent heat are assumed
to be governed by molecular diffusion only, where the transporting eddies are originating
from the outer layers, and thereby enter and leave the interfacial layer randomly. The
exchange between the two sublayers by molecular diffusion, are only active when the
turbulent eddies are in contact with the surface, when a turbulent eddy leaves the surface, the
exchange process stops until a renewal of molecular diffusion can be achieved, by a new eddy
taking its place at the surface. Since the turbulent transport is way more effective than the
molecular diffusion, the interfacial sub layer works as a bottleneck in the transport at the airsea interface.
7.2 COARE (2.0) algorithm
COARE uses a modified bulk algorithm, where it starts by taking the input values of u, T,Ts,
q, R, Rl, Rs and correct Ts and qs from the previous run. Then all neutral transfer coefficients
are given the value of 1.1e-3 and compute the values for , and from eq(2-4). Then the
stability iteration follows: first compute
from:
(62)
Where the boundary layer height is set to =600. Then is determined by an adjusted form
of the Charnock equation (Charnock et al., 1955) for the roughness length, expressed as:
(63)
In the next step the Roughness Reynolds number
is computed by
. The
roughness Reynolds number for temperature and humidity
and
are computed from
23
empirical relations to
, and then
and
can be achieved, and neutral transfer
coefficients determined (Fairall., 1996).
After that comes the steps dealing with stability, first the - functions are computed from the
scalar profile function:
(64)
where
, and
is an empirical constant equal to 12.87. Then the stability
dependent transfer coefficients are determined by first compute 6-9 and then 10-13:
(65)
(66)
(67)
(68)
(69)
(70)
Then
,
and
can be determined according to:
(71)
(72)
(73)
24
The final step is to calculate the Webb-corrected fluxes, and account for effects from
precipitation and cool-skin effect. In the latest COARE 3.0 several steps to improve the model
has been taken including: shortened the stability iteration from 20 to 3, adjustment of the
profile stability functions, redefined transfer coefficient in terms of mixing ratio, both the
velocity and scalar roughness lengths have been changed, for the velocity roughness the
original fixed value of the Charnock parameter has been changed to a one that increases with
wind speed for the winds between 10-18m/s (Fairall et al., 2003).
8. Wintertime Air-Sea fluxes of CO2
The Polar Regions plays an important role, affecting world climate, as a key player for the
global energy circulation, energy is transported between the Inter Tropic Convergence Zone
(ITZ) and the Polar Regions. At latitudes from 50°N and 50°S and above warm surface poleward waters meet and mix with deep cold subpolar waters, rich in nutrients biological activity
is enhanced and a drawdown of
is seen. In addition, wintertime at high latitudes, large
temperature gradients between ocean and atmosphere are common, causing unstable
stratification and deep convection both in air and water enhancing air-sea exchange of CO2.
This together makes the oceans in the high latitudes to act as large sinks of CO2 (Takahashi et
al., 2002).
8.1 Air-Ice -Sea interaction
The ice extent in the Polar Regions has a seasonal cycle, differs from one year to another and
covers at its maximum extent 8% of the world ocean are and is known to have large effect on
the local climate. The high albedo for sea ice, up to 0.9 (Kraus and Businger, 1994), reflects
most of the incoming shortwave radiation, in contrast to open water. Sometimes horizontal
differences in albedo, due to differences in ice thickness leads to formation of “thermal
winds” (Anderson and Neff, 2008).
Fluxes of CO2 between air and sea through brine channels within the ice, formed by rejection
of salt, was first recognized by Goznik (1976). However since the contribution from air-icesea fluxes to the global net flux have not been verified by field observations, the ice is
considered as impermeable layer in most climate models. In the state of the art model
COARE 3.0, the only corrections made, concerns the fetch, macro-structure of the ice and
turbulence induced by snow particles, while air-ice-sea exchange processes such as brine
channels and chemical reactions within the ice are left out.
During the last years, evidence for a direct exchange of CO2 between sea ice and atmosphere
has been presented (Delille, 2006; Semiletov et al., 2004; Zemmelink et al., 2006). In fig. 6 a
schematic picture of today’s knowledge in the seasonal cycle air-ice-sea exchange of CO2 is
presented. In early winter during freezing processes salt is rejected and thereby solved in the
underlying sea water, causes a production of dense cold saline water below the ice which
sinks down to deeper layers. The rejection of salt gives in an increase of the mole fraction of
25
CO2 within the ice and by that also an increase of the partial pressure of CO2, which results in
an oversaturation of CO2 according to eq. 55., This effect from the freezing process results in
two important features, sea ice becomes oversaturated of CO2 resulting in a flux of CO2
through the brine channels from the sea ice to both atmosphere and the under lying water, as
long as the ice is permeable to gas exchange. With time temperature at the air-ice interface
decreases and the brine channels becomes smaller and when the temperature goes below the
threshold of permeability, -7°C to -8°C (Gosnik et al., 1976 ; Golden et al., 1998), the air-ice
gas exchange will theoretically be intermitted. The flux will then be directed from ice into the
sea, where the cold saline CO2 content water sinks down towards deeper layers. However
field observations from the Antarctic sea ice with temperature below -8°C shows that the sea
ice is oversaturated in CO2 and releases CO2 to the atmosphere (Delille, 2010). Recent studies
(Papadmitriou et al, 2004; Delille et al., 2007; Diekmann et al., 2008) have shown that during
ice formation carbonate (
) and Calcium (
) reacts and precipitation of CaCO3 occur
within the ice, with in the process CO2 is rejected out of the ice and that these effect has the
potential to act as a significant sink of atmospheric CO2 (Rysgaard et al., 2007). When
temperature within the air-surface ice is -8°C<T<-5°C the brine channels open and fluxes
from the ice to the atmosphere can occur. In the late ice season when temperature goes above
-5°C the
crystals, formed during sea ice growth, will now during melting be resolved
in water as
and
and thereby consumes CO2, leading to an undersaturation of CO2
within the ice. In addition the springtime biological activity starts within the ice lowering the
CO2 concentration, these together causing a flux of CO2 both from atmosphere and water into
the ice.
Fig 6. Modified from (Delille, 2010). Schematic of our current understanding of CO2
dynamics within sea ice and related air-ice-ocean CO2 exchange during all phases of the ice
growth and decay cycle, with fluxes out of the ice (red) and into the ice (green).
26
Therefore the ice extent plays a crucial role for the parameterization of the air-sea exchange of
CO2, both for the regional and the global distribution. Unfortunately the processes governing
the air-ice-sea gas exchange are not well understood. The lack of alternative of how to relate
the transfer velocity to the flux of gases, the usual wind dependent function is used, developed
for open sea conditions. However these wind dependent functions for the transfer velocity has
an erroneous theoretical basis and shows no empirical evidence to be applied in the presence
of sea ice (McPhee, 1992; Loose and Schlosser, 2010). In regions covered by sea ice,
turbulence can be produced through buoyant convection and by promotion of a current shear
between ice and water (McPhee, 1992: Morison et al., 1992). In a partly ice covered surface
steep short-fetch wind-generated waves are built up, that potentially would increase the gas
transfer velocities. Adding these effects the gas transfer velocity has the potential to be above
what would be expected for the open ocean (Loose and Schlosser, 2011).
The usual theory states that the ice cover insulates the ocean from the atmosphere during mid
winter and thereby almost completely reduces the exchange of heat, moist and gases between
the ocean and atmosphere. However according to Loose and Schlosser (2011), the ice cover
instead of trapping CO2 beneath the ice until springtime, it delays the gas fluxes until
springtime and ice melt. Instead of the wind dependent function, Looser and Schlosser (2010)
used a trace gas mass balance function. By using data from Ice station Weddell they showed
for the Southern Ocean south of 50°S, that the net CO2 flux through sea ice cover represents
14-46% of the net annual air-sea flux, depending on which relationship between sea ice cover
and k660 that is used. In average the k660 across the air-ice-sea interface with 100% ice cover
were 0.11m d-1.
The model they used also showed that 68% of net annual CO2 flux in the sea-ice zone occurs
during springtime within the marginal ice zone, thus they emphasis for a more accurate
parameterization of the gas air-ice-sea flux. In prior studies larger values of k660 has been
reported, Fanning and Torres (1991) found k660 to be in the range 1.44m d-1 - 3.36m d-1 from
what they described as almost complete ice cover, and for a ice cover less than 70% a k660 of
2.14 m d-1 – 3.36 m d-1was found.
Recent field measurements of CO2 (Sörensen, 2010) using the eddy covariance technique,
over a Greenland fjord with a thickness of ice of 0.5-1m, showed on average small downward
fluxes of CO2. However occasionally larger upward fluxes were found, at the same time high
upward fluxes of latent heat, a weak downward fluxes of sensible heat occurred, for a
thickness of ice of 0.5-1m. During late winter, Papakyriaiou and Miller (2011) reported a
maximum hourly efflux of CO2 of 1.0 µmol m-2 s-1 and -3.0 µmol m-2 s-1 into the ice, on
average a efflux of 0.36 µmol m-2 s-1 was found. These fluxes are much greater than former
studies of CO2 fluxes over ice and are comparable with CO2 fluxes found for open sea
conditions.
The question on how to relate ice cover to the gas transfer velocity remains unsolved.
Takahashi et al. (2009), suggested a linear dependence (fig. 5) that is used in the latest climate
27
models, however this linear dependence on wind speed does not match with the only few
good prior field estimates of k660 to the ice cover, Faning and Torres (1991) and from Ice
station Weddell (1992). A recent laboratory study (Loose et al., 2009), where two different
trace gases SF6 and O2 were studied for fractions of ice/open water, disagree with the linear
dependence of T09. In the study measurements from Fanning and Torres (1991), Loose and
Schlosser (2010) where used, and they were all clustered above the linear 1:1 dependence
between k/k100% open water and fraction of open water (f), demonstrating that the gas transfer do
not scale linearly with the fraction of open water. This since the turbulence dissipation
beneath does not have to be a strict function of fetch, thus other processes such as buoyant
convection and shear stress affects turbulent regime in the air-ice-sea exchange ass discussed
above. Papakyriaiou and Miller (2011) found a relationship between the flux of CO2 and wind
speed, where fluxes into the ice where associated with warming and high wind speeds, while
effluxes in general occurred for wind speeds <4.5m s-1. This suggests that along with
temperature, turbulent exchange with, and ventilation of, snow might be an important feature
of the exchange process.
9. The Baltic Sea and the site of Östergarnsholm
The Baltic Sea is a brackish inland sea located in northern Europe, from 53°N to 66°N latitude
and from 20°E to 26°E longitude and regarded as a coastal sea. The basin of the Baltic Sea is
formed by glacial erosion during the last few ice ages. The Baltic Sea has a net precipitation
of about 1500 m3s-1 and a river input of 15,000 m3s-1 (Bergström and Carlsson, 1994). With
the large amount of fresh water follows a large amount of nutritients, organic and in organic
carbon. This creates a unique dynamic system with horizontal and vertical gradients of
variables controlling the efficiency of the air-sea exchange of carbon dioxide, such as
temperature, salinity, pH and alkalinity (Omstedt et al., 2004).
In the recent years much attention has been paid to the coastal seas regarding the air-sea CO2
system. Even though shelf regions and small shallow seas occupy a marginal portion of the
world´s total ocean surface, their role in the global marine primary production are of major
importance. The biological production constitutes a key parameter controlling the state of the
marine CO2 system (e.g., in the Baltic sea; Thomas and Sneider, 1999), thus plays an
important role in the CO2 cycle. The sea can act as source or a sink of carbon dioxide, which
it, mainly depends whether the net biological production exceeds the mineralization. During
periods when the biological production exceeds mineralization, the sea more likely acts as a
sink of CO2, and CO2 is taken up by the sea from the air. Roughly one can say that the cold
productive water in high latitudes acts as a sink of CO2 and the upwelling regions as a source
(Takahashi et al., 2002). Studies regarding the sink/source distribution over the European
shelf regions (e.g., Borges et al.2006; Chen and Borges, 2009) finds the continental shelfes
are sinks while the analyzed estuaries are sources of atmospheric CO2.
28
9.1 The Östergarnsholm site
In the Baltic Sea measurements has been done semi-continously since 1995 in a tower, on the
southern tip of the island of Östergarnsholm (57°27’N, 18°59’E) east of Gotland.
Temperature, wind speed and wind direction are measured at five levels 7, 11.5, 14, 20 and 28
m, and the relative humidity at 8 m. At 9, 16.5 and 25 m the wind components are measured
with high frequency by two sonic anemometers, Windmaster (9m, 25m) and SOLENT
1012R2 (16.5m) both from Gill Instruments, Lymington, UK. The CO2 and humidity
fluctuations are measured with an semi-enclosed Licor 7200 analyzer at 9 m and with a Licor
7500 open path gas analyzer at 25 m both (LICOR-Inc.m Lincon, NE, USA). Atmospheric
CO2 is also measured every half hour with an infrared gas analyzer (IRGA), PP-systems.
Since 2005 a SAMI sensor measuring
and SST in water attached in a buoy 1 km SE of
the island (fig 7) has been running.
Figure 7. Modified from Rutgersson et al. (2008). Upper figure: The Baltic Sea and the site of
Östergarnsholm (red dot). In the lower figure: The position of the tower and the Sami sensor,
are marked by arrows. The thin line in the lower figure displays isolines of water depth
Measurements made at Östergarnsholm have been shown to represent open sea conditions for
wind direction 80-210°, (Högström et al., 2008), however for CO2 open sea conditions is find
for wind direction 80-160° (Rutgersson et al., 2008), where the footprint covers the place area
29
of the buoy. Using measurements for the period 2005-2007 Rutgersson et al.( 2009) found
that the Baltic Sea shows a significant seasonal cycle of CO2 with peak values above 800
µatm in wintertime and as low as 150 µatm in summer.
10. MicroTX3 oxygen sensor
The fiber optical microsensor, Microx TX3, measures the concentration fluctuations of
oxygen. The instrument uses fiberoptic technique, were it first excites an indicator molecule
and then detects the decrease of luminescence of the indicator luminescent molecule in the
presence of oxygen. First the indicator molecule gets excited by a sinusoidal light, in the
absence of oxygen the indicator molecule emits light, as it returns to ground state, which then
is detected by the instrument. In the presence of oxygen a collision with an oxygen molecule
occur, excited energy is then transferred from the indicator molecule to the oxygen molecule.
The oxygen molecule gets excited and eventually returns to ground state and thereby emits
light, not detected by the instrument.
The relation between the concentration of oxygen and the intensity of the luminescence is:
I0
 1  K SV O2 
I1
(74)
where I1 is the intensity of the luminescence if oxygen is present, I0 is the intensity of the
luminescence if no oxygen is present, Ksv is the Stern-Volmer constant and [O2] is the oxygen
concentration. With this relation, the concentration of oxygen can be calculated from the
luminescence. The Stern-Volmar constant is the product between the quenching constant, kq,
and the lifetime of the excited state of the indicator molecule, if no oxygen is present. The
quenching constant in turn has to be determined experimentally.
It is difficult to directly measure the luminescence. Therefore the luminescence lifetime is
used as measure of the oxygen dependent quenching. The advantage to measure the decay
time ( ) instead of the luminescence is that the decay time does not depend on the sensitivity
of the detector, it is independent of the concentration of the indicator molecule and it is not
influenced by optical properties of the sample.
A technique called phase modulation is used (fig 8a) to measure the decay time of the
luminescence from the indicator molecules. The decay time is defined as the time between the
excited signal and the molecule emitted signal. In the presence of oxygen the decay time (τ1)
is shorter compared to the decay time in oxygen free air (τ0). Hence the time delay between
excited and emitted signal can be represented as a phase angle (Φ). The presence of oxygen in
the sample causes a shift in the phase angle. In figure 8b the phase angle shift is represented.
When there is no oxygen in the sample the phase angle is Φ0, and when there is oxygen in the
sample the phase is shifted, Φ1.
30
a)
b)
Fig 9. A. (Huber and Krause, 2006) Decay time in the presence of oxygen ( ) and decay time
in the absence of oxygen ( ). b) Phase angle shift were (Φ0) denotes the phase angle
between reference signal and measuring signal in oxygen free air, (Φ1) denotes the phase
angle between reference signal and measuring signal in air containing oxygen. (From Huber
and Krause, 2006)
Using the concept of phase angle, eq. 1 can be rewritten as:
I0

tan  0
 1  K SV O2   0 
I1
 1 tan 1
(75)
The oxygen concentration can thus be calculated from the phase angle shift instead of
luminescence. The phase angle related to the oxygen concentration can be seen in figure 9.
Fig 9. (Huber and Krause, 2006). Phase angle [°] related to the oxygen concentration in air
with saturated water vapor [%], relative humidity=100%.
31
For eddy covariance measurements a high sampling frequency is needed. The Microx TX3,
has a response time (which implies that 90% of the real oxygen concentration has been
reached) of up to 0.5 s and can measure fluctuations up to 1 Hz. The instrument is set to give
the amount of O2 as % air-saturation, where 100 % air-saturation is equal to the normal
volume content of oxygen of 20.95% at the pressure 1013.25hPa.
The unit of % air-saturation can be converted into µmol/l through the following equation:
1
 mol  p atm  p w T  %air  saturation 
CO2 

0.2095 T 

pN
100
vM
 l 
(76)
where patm is the atmospheric pressure, pN the standard pressure (1013hPa), pw(T) is the water
vapor pressure of saturated air at temperature T (K), VM is the molar volume (22.414 l/mol)
and α(T) is the Bunsen absorption coefficient at temperature T. The Bunsen absorption
coefficient is calculated using the following equation:
ln 10 3  T  
8.553  10 3
 23.78 ln T  160.8
T
(77)
11. Future work
Recently a high frequency instrument for long term measurements of oxygen has been installed in the
tower at the site of Östergarnsholm. From these new measurements combined with the existing
performed at the tower and the buoy, fluxes and transfer velocity of oxygen can be computed and
compared with those for CO2 in order to refine the relation between transfer velocity and the processes
affecting it. In addition the transfer velocity of oxygen will add new knowledge to the role of the ocean
in the global carbon cycle. Also a newly developed enclosed Licor (7200), capable of measuring CO2
concentration during rainfall has been installed. Then fluxes of CO2 can be computed by the Eddycovariance method, and the effects from rainfall on the transfer velocity to be studied. The recent
findings of a feasible direct CO2 gas exchange between sea ice and the atmosphere opens up a new
field for research in air-sea interaction at high latitudes. Since only very few field studies of CO2
fluxes over ice with high frequency instruments has been made, there is a large lack of knowledge
concerning the processes affecting the transfer velocity, thus more field studies in different
environments has to be performed.
32
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