2003 in the European Area showed an unusual situation: soil humidity was very low in most of the
territory, because of the lack of precipitations, and soil temperature was very high in north-western
Italy. High temperature and low humidity are decisive factors for net radiation, higher than normal,
which consequently implies fluxes of sensible and latent heat to be higher than normal. In particu-
lar, because of the low soil humidity, the latent heat flux has remained equal to or lower than the
standard, so the sensible heat flux has been much higher than normal. This has caused excessive
heating of the lower atmosphere, further decreasing evapotranspiration. This positive feedback has
favoured the existence of anticyclonic conditions, persisting the whole trimester.
This example shows that it’s necessary to know the values of soil humidity and turbulent heat
fluxes in order to explain meteorological or climatologic phenomena. Nevertheless there are only a
few sensors to measure soil humidity and turbulent heat fluxes in the world. Moreover soil humidity
is variable and it can have different values within the same soil. Satellites could be used, but many
of them see atmospheric water, while the deduction of surface humidity content from specific satel-
lites is still at a preliminary phase, and anyway it doesn’t provide the measures of turbulent fluxes.
That’s why numerical models are used to provide data groups of soil humidity and temperature and
of sensible and latent heat fluxes.
My thesis work is developed within a collaboration of a ASEM and MAE (the Foreign Office) in-
ternational exchange project, and the purpose of my thesis has been to determine, evaluate and ana-
lyse the typical trends of the energetic (turbulent fluxes of sensible and latent heat) and hydrologic
(precipitation, evapotranspiration, runoff and drainage) balances in the surface layer. Since the
number of employable direct measures is fragmentary and exiguous, and also insufficiently flexible
for the climatologic analysis, I’ve used the CLIPS (surface parameters climatology) methodology,
already used in Europe in the past: one takes data from meteorological stations, creating datasets
with continuous values by data interpolation, and one runs a reliable SVAT model, to compute the
parameters in the surface layer. In my thesis work I’ve used LSPM (Land Surface Process Model,
Cassardo et al., 1995), one-dimensional diagnostic SVAT model, which treats with particular regard
the estimation of soil and vegetation characteristic parameters and the computation of soil tempera-
ture and humidity, and I’ve focused on the summer monsoon period in South Korea. At the end of
the simulations, I’ve checked some parameters, computed by LSPM, by comparison with available
experimental measures. I’ve made graphs with the trends of the daily average values and I’ve statis-
tically evaluated the differences.
This work is important for the “Amedeo Avogadro” General Physics Department of Turin, because
I’ve applied an Italian model created in Turin using the typical data of a monsoonal climate, as well
8
as for South Korea, because nobody had ever worked on this topic and up to now there weren’t any
available maps such as the ones I’ve created.
The parts of my thesis can be divided into 4 blocks:
1) The first part describes the features and the governing equations for the atmosphere and for the
atmospheric boundary layer.
2) The second part describes the LSPM model.
3) The third section is dedicated to the operations of data preparation (pre-processing): in particular
the operations of densifying, correction and interpolation of data in order to make them fit to be
read by the LSPM model.
4) The fourth and last part is dedicated to the use of the simulation model, to the comment of the
two-dimensional maps on the Korean territory of the monthly average values of some key parame-
ters and of the graphs of their average daily trends.
First of all, in particular, I’ve deemed it appropriate to introduce a few concepts on the atmosphere
physics and the main equations; in the second chapter I’ve introduced the boundary layer, with its
main features, since this is the atmospheric layer where the variables studied in my thesis have been
measured; then I’ve examined the turbulent motion equations in the third chapter, with particular
regard to the stability parameters and to the similarity theory of Monin-Obukhov. In the fourth
chapter I’ve inserted a section concerning the hydrologic and radiative balance at the earth surface,
and a part concerning the soil subsurface layer and its features, analysing also the hydraulic and
thermic conduction within the soil.
In the fifth chapter I’ve introduced the theory of the LSPM model. I’ve particularly lingered over
the radiating fluxes of linear momentum and of sensible and latent heat, distinguishing between
vegetated and non-vegetated surfaces. The final topics of the chapter are the water cycle, the analy-
sis of the hydrologic balance for the different soil layers and the heat transfer in the soil.
The sixth and seventh chapter concern, respectively, the preliminary treatment of data from simula-
tions in South Korea, and the verification of the model reliability, especially with the graphical and
statistical analysis of observed and simulated data. In the seventh chapter I’ve also provided and
commented the monthly average maps of some significant variables (solar and net radiation, sensi-
ble and latent heat fluxes, soil surface temperature and humidity) on the Korean territory.
9
Chapter 1
THE ATMOSPHERE
1.1 INTRODUCTION
The Earth’s atmosphere is a mixture of gases, mainly made out of nitrogen and oxygen, whose
properties are defined up to some hundreds kilometres of altitude, and to which the principles and
procedures of physics are applied.
Going from Earth to space, the interplanetary explorations have observed there’s no discontinuity in
matter distribution, so the Earth’s atmosphere appears as a local thickening of the interplanetary at-
mosphere rather than a peculiar element of Earth.
In order to analyse it from different points of view, the atmosphere is often divided according to
classifications related to some physical properties: the most used among them is related to the tem-
perature trend as a function of altitude, showed in table 1.1, where the names of the atmospheric
layers, their average thickness and the temperature trend are showed. This is an average temperature
in time and space.
Figure 1.1 – Temperature trend as a function of altitude in the atmosphere
Let’s briefly see the general nature of these layers.
10
1.1.1 Troposphere
The main feature of the troposphere is the temperature decrease with height, whose rate (dT/dz) is
called “lapse rate”. The mean value of the lapse rate in the troposphere is about 0.65 °C/100 m, with
deviations from the average seasonal values which can reach ρ 0.3 °C/100 m in a given place. The
troposphere is the seat of the fogs, of the main types of clouds and of the storm activities; moreover,
most of the atmospheric mass is concentrated within it (from 75% at medium and high latitudes, up
to 90% at low latitudes). Furthermore the troposphere can be divided into:
1- lower troposphere, or boundary layer, which will be described later on, extending from the
Earth’s surface up to about 1 ÷ 1.5 km;
2- medium troposphere, from 1÷1.5 up to 6÷7 km;
3- upper troposphere, from 6÷7 km up to the tropopause.
11
Name of the layer Lower boundary [km] Upper boundary [km] Trend of T(z)
Troposphere 0 10 decreasing
Tropopause 10 30 constant
Stratosphere 30 48 increasing
Stratopause 48 53 constant
Mesosphere 53 75 decreasing
Mesopause 75 90 constant
Thermosphere 90 120 increasing
Thermopause 120 800 -
Exosphere 800 infinite -
Table 1.1 – Division of the atmosphere according to its thermic features
The wind regime is characterised, near the ground surface, by a dramatic reduction, mainly due to
friction, of the speed modulus, which goes to zero at the ground and in a thin layer adjacent to it.
Above this layer the wind speed varies, within the boundary layer (BL), at first quickly, then more
slowly (logarithmic profile). A remarkable feature, connected to the increase of the wind speed
modulus along with height, is the corresponding direction change (Elkmann’s spiral). Above the
BL, the wind changes with altitude are mainly due to the horizontal temperature gradients; since the
highest temperatures are observed in the equatorial and subtropical regions, whereas the lowest ones
are observed in the polar regions, the influence of the horizontal temperature gradient causes a
strengthening of the wind occidental component and a weakening of its oriental component with
height; as the occidental streams prevail over the oriental ones above the BL at medium latitudes,
the result is that wind speed increases along with altitude.
Since there are remarkable horizontal temperature differences in the air masses, as it happens near
evident frontal zones, the so called “jet streams” develop in the troposphere and in the stratosphere;
they’re zones of strong winds, relatively limited transversally and very extended longitudinally; the
higher is the temperature difference between the air masses, where the jet stream develops in the
frontal transition zone, the greater is the wind speed along the jet axis.
The troposphere is the place of life: all the plants and the living beings live inside it, using some of
the gases it’s made out of. Since the first launches of sounding balloons at the beginning of the XX
century it had been observed that the temperature decrease along with height ended at a certain
12
point; at first the decrease became slower, then the temperature became stable, with an isothermal
distribution. This feature has been confirmed in all the following radio-soundings: the atmospheric
layers characterised by a null or negative average lapse rate (temperature increasing with height)
constitute the stratosphere, extending more or less from 11 up to 50 km. The transition layer be-
tween the troposphere and the stratosphere, whose thickness ranges between a few hundreds meters
and 1-1.5 km, is called tropopause; within it the lapse rate varies from 0.6-0.8/100 m in the lower
part, to zero, or even to negative values, in the upper part.
1.1.2 Stratosphere
It’s the atmospheric layer above the tropopause, and it extends up to an altitude of 50-60 km. Here a
phenomenon called thermic inversion takes place: whereas, in the troposphere, the temperature de-
creases along with height, in the stratosphere it increases, until it reaches nearly 0 °C. This is due to
the existence of an ozone layer, the ozonosphere, which absorbs most of the ultraviolet solar radia-
tions. In the stratosphere the components are more and more rarefied, water vapour and atmospheric
dust decrease. The temperature of the transition layer, the stratopause (at an altitude of about 45-55
km), is close to 0 °C, with deviations of ρ 20 °C (data coming from rockets).
1.1.3 Mesosphere
In this zone, extending from 50 to 80 km of height, the atmosphere isn’t any more influenced by the
Earth’s surface, and it’s constant at all latitudes; the temperature meanly decreases with altitude.
The temperature near the top of the mesosphere (85-90 km) is around -80 °C in summer , at high
and medium latitudes, and -40 °C in winter; here is the transition layer called mesopause. Beyond
the mesopause, at the height of around 100 km, the air is so rarefied that it doesn’t offer any resis-
tance to the bodies’ motion, so it becomes possible to move with the orbital motion. That’s why in
astronautics the mesosphere is considered the border with space.
1.1.4 Thermosphere
Beyond the mesopause the thermosphere begins. The gases existing in this layer are so rarefied that
it’s more appropriate to talk about a plasma made up of atoms, ions and molecules, which receive in
certain bands the direct solar radiation and consequently are mainly in the ionized state (the thermo-
sphere, with the upper layers of the mesosphere, constitutes the Earth’s ionosphere). The tempera-
ture, to be intended in a kinetic sense as a measure of the average kinetic energy of the plasma par-
ticles, rises along with altitude in this layer, because of solar irradiation, and it reaches 1700 ºC at
13
its outer boundary. At the border between mesosphere and troposphere, the auroras borealis take
place.
1.1.5 Exosphere
It’s the outmost part of the Earth’s atmosphere, where the chemical composition radically changes.
The exosphere has no real upper boundary, and its components are mainly hydrogen and helium. By
means of indirect observations it’s been inferred that the (kinetic) temperature in the exosphere in-
creases with height, reaching or even passing 2000 ° C.
Another classification can be made according to the influence of the Earth’s surface processes on
the atmosphere. In this case the planetary boundary layer (0-1500 m approximately) is defined as
the part of lower troposphere where the phenomena of energetic and hydrologic exchange between
soil and atmosphere are fundamental, and where the human presence has a greater influence.
This is the layer I’ve studied in this thesis, and I’m going discuss it in detail afterwards. The remain-
ing part of the troposphere is simply referred to as free atmosphere (over 1500 m), that is not inter-
ested by surface friction forces.
1.1.6 Atmospheric regions
From a theoretical point of view, the description of the atmospheric motions is very complex. For
this reason a further classification of the atmosphere is necessary. It considers the spatial-temporal
orders of magnitude of the different physical processes and it makes it possible to simplify the em-
ployed equations, helping their modelling and theoretical study. In fact the motion scales express
the distance and the time interval during in which an appreciable variation of a certain physical
quantity is observed. In the atmospheric case, the phenomena have scales ranging from molecular to
planetary dimensions, from the fractions of second to several days. Accounting for all of this, three
atmospheric regions are usually referred to:
i) Microscale region: it’s characterised by phenomena taking place in spaces whose linear dimen-
sions range from a few centimetres to some kilometres, in scale times ranging from one second to
some hours. The principal physical processes are those of the earth-atmosphere interface (evapora-
tion, evapotranspiration, turbulent diffusion and convection), which characterise the planetary
boundary layer.
ii) Mesoscale region: the typical phenomena involve distances of a few tens of kilometres and time
periods ranging from a few hours to several days.
14
iii) Macroscale region: it’s characterised by phenomena whose dimensions can be those of the con-
tinents, for instance continental (synoptic) and planetary scale motions, particularly due to the non-
uniform heating of the Earth by the sun. These movements, which in turn are responsible for the
great thermic energy transfers from the equator to the pole, are influenced by the Coriolis force due
to the Earth rotation, by the irregular distribution of the continents on the Earth’s surface and by the
interactions with the great barriers represented by the main relieves.
The scale analysis allows to face in a more direct and simpler way the complexity of the atmos-
pheric dynamics equations, favouring the introduction of simplifications for the physical phenom-
ena that, according to the scale, aren’t meaningful in the study of the atmosphere. For instance, the
Coriolis force, due to the Earth’s rotation, in negligible at the microscale, whereas it’s fundamental
at the macroscale; in the same way the turbulent phenomena are decisive at the microscale and less
important at the macroscale.
In order to decide the influence of every single process, a technique called scale analysis is em-
ployed, that allows to understand which processes can be neglected and which are of great impor-
tance in the considered physical field. This system is aimed at speeding up the computations to be
done, in order to make the obtained results more meaningful, leaving out all the situations consid-
ered irrelevant, and it also allows the simplification of some non-linear terms which would make it
impossible to solve the set of equations governing the system.
By means of restrictive hypothesis, a model of the system is set up, whose correctness can be
checked by comparison between the obtained results and the real observed data. Of course, because
of the mentioned restrictive hypothesis, it’s extremely counterproductive to use a model outside its
validity field. In scale analysis the importance of the considered variables, the amplitude of their
fluctuations and the characteristic spatial and temporal scale of these fluctuation are accounted for.
The scale analysis technique is usually faced by means of the Reynolds’ approach, which provides
for the division of the variables into an average value plus a fluctuation, where the mean can be de-
fined in different ways (temporal, spatial, overall). By this method the non-linear terms representing
the turbulent phenomena are parameterized in a simplified way, starting from the fundamental con-
cepts and considering experimental results as well. The level of description of the turbulence de-
fines the capability of the model to analyse the motion field and the mechanical and thermic turbu-
lent fluxes.
15
1.2 MAIN ATMOSPHERIC QUANTITIES
Before going deep into the analysis of the dynamic processes of the atmosphere and of its governing
equations, I’ll linger over the main atmospheric variables, which are usually more frequently meas-
ured or computed. Here follows a short description of them.
Temperature
Measurement unit: degree (defined as the hundredth part of the interval between water’s boiling and
freezing point, at the atmospheric pressure). It’s measured by the Kelvin (°K) or the Celsius (°C)
scale, with 0 °C = 273.15 °K, or by the Fahrenheit (°F) scale.
Pressure
Measurement unit: Pascal (pressure exerted by a force of 1 N on a surface of 1 m
2
). Usually in me-
teorology the hPa (= 100 Pa) is used. The standard atmospheric pressure of air at sea level is
1013.25 hPa.
Vapour pressure
It’s the partial pressure exerted by the saturated water vapour, it’s noted as “E” (unlike the unsatu-
rated water vapour pressured, noted as “e”) and it’s computed by the empirical Tetens-Magnus for-
mula:
tb
ta
EtE
10)(
0
(1.)
with a = 7.5 and b = 237.3 for water, a = 9.321 and b = 261.24 for ice, and where the temperature t
is expressed in °C. E
0
= 6.1 hPa = saturated water vapour pressure at 0 °C.
Specific humidity
Ratio between the water vapour mass contained in an air mass and the air mass itself: q = m
v
/m ; it’s
usually measured in [g vapour /Kg vapour].
Using the state equation for humid air: pV = mR
d
T(1+0.606s) (1.2)
for dry air: p
d
V = m
d
R
d
T ,
and for water vapour: eV
v
= m
v
R
v
T ,
one obtains, for the specific humidity: q = m
v
/m ≈
ep
e
378.0
622.0
(1.3)
where R
d
= 287.04 J/(Kg °K) = specific constant for dry air, and R
v
= 461.5 J/(Kg °K) = specific
constant for water vapour.
Relative humidity
Percentage ratio between the water vapour mass contained in a certain air volume and the mass that
would be contained at the same temperature in saturated conditions:
16
Ricevi informazioni sui nostri servizi, sulle offerte e non perdere news e consigli su università e lavoro.
