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CHAPTER –I
INTRODUCTION TO OZONE
1.1 Introduction
In the recent past climate change and environmental problems are the major
global issues. The green house gases loading, aerosol loading, ozone depletion,
increasing surface ozone, acid rains are the major environmental problems to the
scientific society. Ozone acts as a beneficial one or harmful depends on the
concentration where it is located. Recently, researchers showed that stratospheric
ozone level has been decreasing in significant level, at the same time due to industrial
and transport activities surface ozone has been increasing. Increasing ozone level
changes the climate through both direct and indirect routes. Since ozone is involved
in climate changes, it is necessary to study the ozone and its measurement in the
present time. Ozone is a secondary photochemical pollutant [1, 2]. Elevated levels
cause health problems, premature deaths, reduced agricultural crop yields, changes in
ecosystem species composition and damage to physical infrastructure and cultural
heritage.
1.2 Ozone and its structure
Ozone is a minor constituent of the atmosphere. It is a form of oxygen. It is an
odorless, colorless gas. At the molecular level, ozone is made up of three oxygen
atoms (O3) with one double bond and one single bond. Normal oxygen, which we
breathe, has two oxygen atoms and is colorless and odorless. Ozone, or triatomic
oxygen, is much more unstable than diatomic oxygen found in air and is therefore a
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strong oxidizing agent. Ozone was discovered by C.F.Schönbein in the middle of the
last century; he also was first to detect ozone in air [3, 4]. Ozone is much less
common than normal oxygen. Out of each 10 million air molecules, about 2 million
are normal oxygen, but only 3 are ozone. Ozone exits in the atmosphere at two
heights. They are starostocpheric ozone and tropospheric ozone. Fig 1.1 shows
structure of ozone.
Fig 1.1- Ozone Structure
1.3 Ozone in Atmosphere
Ozone is important in two layers of the atmosphere. Most ozone (about 90%)
resides in a layer between approximately 10 and 50 km above the Earth’s surface, in
the region of the atmosphere called the stratosphere. Roughly 10% of the Earth's
ozone is found in the lower region of the atmosphere, in the region of the atmosphere
called, troposphere. Most ozone is produced naturally in the upper atmosphere or
stratosphere. While ozone can be found through the entire atmosphere, the greatest
concentration occurs at an altitude of about 25 km [5]. This band of ozone-rich air is
known as the "ozone layer". The stratosphere ranges from an altitude of 10km to
50km and it lies above the troposphere (or lower atmosphere). Even though both
types of ozone are exactly the same molecule, their presence in different parts of the
atmosphere has very different consequences. Stratospheric ozone blocks harmful
solar radiation - all life on Earth has adapted to this filtered solar radiation.
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Stratospheric ozone shields us from the harmful effects of the Sun's ultra-violet
radiation. Fig 1.2 shows ozone in atmosphere.
Fig 1.2 - Ozone in atmosphere
1.4 Why is the ozone layer important?
Ozone's unique physical properties allow the ozone layer to act as our planet's
sunscreen, providing an invisible filter to help protect all life forms from the Sun's
damaging ultraviolet (UV) rays. Most incoming UV radiation is absorbed by ozone
and prevented from reaching the Earth's surface. Without the protective effect of
ozone, life on Earth would not have evolved in the way it has.
1.5 Ozone measurement units
Various units are used to measure surface ozone. The most common surface
ozone measurement is ppm, and is used to measure ozone in air and ozone dissolved
into water.
(i) ppm = parts per million is a measurement of concentration. If we state there is 1
ppm ozone that means for every 1 million parts of gas, 1 of these is ozone.
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(ii) ppb = parts per billion. For example 0.1 ppm = 100 ppb.
(iii) mg/l = milligrams of ozone per liter. This indicates how many milligrams of
ozone there are in one liter of total volume. Mg/l can be used to indicate the
concentration of ozone in gas or liquid. 1 mg/l of ozone = 1 ppm of ozone in
water. Due to the density of air this is no longer true and 1 ppm of ozone = 2140
mg/l. This is a common term used to measure the amount of ozone dissolved into
water.
(iv) µg/ml = microgram per milliliter. This indicates how many micrograms of ozone
there are in one milliliter of total volume. µg/ml can be used to indicate the
concentration of ozone in gas or liquid. 1 µg/ml = 1 mg/l = 1g/m3 = 1 gamma.
(v) G/m3 = Grams of ozone per Cubic Meter is a measurement of concentration.
This indicates how many grams of ozone there are in one cubic meter of total
volume. This can indicate volume of a gas or liquid. g/m3
is most commonly
used to measure the concentration of ozone in a gas stream. 1 g/m3
= 1 mg/l = 1
ppm of ozone in water 1 g/m3
= 467 ppm of ozone in air.
(vi) %by weight (Percent by weight) is a measurement of concentration. This
indicates the percentage (%) of ozone within a given gas stream. This is a very
common method to illustrate the concentration of ozone from an Ozone
Generator. This is more complicated than g/m3
as the weight of gas changes if it
is air, or oxygen gas that the ozone gas is mixed with 1% Ozone = 12.8
g/m3
Ozone in air 1% Ozone = 14.3 g/m3
Ozone in oxygen.
(vii) G/hr (gm/hr) = grams of ozone per hour is a measurement of ozone production.
This is the most common method of measuring the output of an Ozone
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Generator. We can measure the concentration of ozone in g/m3, then when we
calculate for flow rate with a measurement such as LPM (liters per minute), we
can determine how many grams of ozone are produced in one hour of time.
(viii) Mg/hr = milligrams of ozone per hour. This indicates the same thing as g/hr
only on a smaller scale. Smaller Ozone Generators may be rated in mg/hr.1 g/hr
of ozone = 1,000 mg/hr of ozone production.
(ix) Lb/day = Pounds per Day is a measurement of ozone production. This is a
common measurement of ozone production for large Ozone Generators, and is
commonly used in some industries within the ozone world Lb/day of ozone =
18.89 g/hr ozone production.
1.6 The abundance and measurement of ozone levels
The abundance of ozone in the atmosphere is measured by a variety of
techniques as shown in figure 1.3. The techniques make use of ozone’s unique optical
and chemical properties. There are two principal categories of measurement
techniques: local and remote. Ozone measurements by these techniques have been
essential in monitoring changes in the ozone layer and in developing our
understanding of the processes that control ozone abundances.
1.6.1 Local measurements
Local measurements of atmospheric ozone abundance are those that require
air to be drawn directly into an instrument. Once inside an instrument, ozone can be
measured by its absorption of ultraviolet (UV) light or by the electrical current
produced in an ozone chemical reaction. The latter approach is used in the
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construction of “ozonesondes,” which are light weight. Ozone measuring modules are
suitable for launching on small balloons. The balloons ascend far enough in the
atmosphere to measure ozone in the stratospheric ozone layer. Local ozone-
measuring instruments using optical or chemical detection schemes are also used
routinely on board research aircraft to measure the distribution of ozone in the
troposphere and lower stratosphere.
1.6.2. Remote measurements
Remote measurements of ozone abundance are obtained by detecting the
presence of ozone at large distances away from the instrument. Most remote
measurements of ozone rely on its unique absorption of UV radiation. Sources of UV
radiation that can be used are the Sun and lasers. For example, satellites use the
absorption of UV sunlight by the atmosphere or the absorption of sunlight scattered
from the surface of Earth to measure ozone over nearly the entire globe on a daily
basis. A network of ground-based detectors measures ozone by the amount of the
Sun’s UV light that reaches Earth’s surface. Other instruments measure ozone using
its absorption of infrared or visible radiation or its emission of microwave or infrared
radiation. Total ozone amounts and the altitude distribution of ozone can be obtained
with remote measurement techniques. Lasers are routinely deployed at ground sites or
on board aircraft to detect ozone over a distance of many kilometers along the laser
light path.
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Fig 1.3-Ozone measured by various instruments techniques
1.7 The Formation of Stratospheric Ozone
Ozone is created in the stratosphere when highly energetic solar radiation
strikes molecules of oxygen (O2) and cause the two oxygen atoms to split apart. If a
freed atom bumps into another O2, it joins up, forming ozone (O3). This process is
known as photolysis. Fig1.4 refers ozone formation in stratosphere. Ozone is also
naturally broken down in the stratosphere by sunlight and by a chemical reaction with
various compounds containing nitrogen, hydrogen and chlorine. These chemicals all
occur naturally in the atmosphere in very small amounts. In an unpolluted atmosphere
there is a balance between the amount of ozone being produced and the amount of
ozone being destroyed. As a result, the total concentration of ozone in the
stratosphere remains relatively constant. At different temperatures and pressures (i.e.
varying altitudes within the stratosphere), there are different formation and
destruction reaction rates. Thus, the amount of ozone within the stratosphere varies
according to altitude. Ozone concentrations are highest between 19 and 23 km.
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Fig 1.4–Stratospheric ozone production
1.8 Effect of climate change on stratospheric ozone
Climate change affects stratospheric ozone in two principal ways. The first is
through in-situ hanges within the stratosphere, from radiative-chemical mechanisms,
associated with the changes in anthropogenic greenhouse gases (for the stratosphere,
this is mainly from the changes in carbon dioxide).The second is through changes in
stratospheric wave forcing and hence in the Brewer-Dobson circulation which result
from changes in tropospheric climate. That the latter has the potential to be a
significant contributor is evident from the contrast between the two hemispheres in
the current climate, which results from differences in wave forcing. Whilst in the real
atmosphere these two processes occur together and cannot be isolated, in a model
they can be. In doubled-CO2 experiments with chemistry-climate models, [6]
separated the two processes by doubling CO2 separately in the troposphere and the
stratosphere, while [7] and [8] instead did so by separately doubling CO2 and making
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the corresponding changes to sea surface temperatures and the associated sea ice
distributions.
Climate change may cause an increase in temperature in the troposphere and a
decrease in temperature in the stratosphere. This decrease may delay the recovery of
the ozone layer at the Arctic and Antarctic by several years, due to an increase in
clouds in the stratosphere. The greenhouse gases methane and nitrous oxide may also
affect stratospheric ozone by chemical interactions. This may have a positive or
negative effect [9]. The magnitude of this interaction is uncertain and depends on the
emissions of methane and nitrous oxide and the chlorine concentration. Aircraft emit
nitrogen oxides (NOx) directly in the upper troposphere and lower stratosphere. This
is the only anthropogenic source of NOx in these regions. The impact of the current
fleet of aircraft on the observed ozone depletion is unknown but probably small. The
effect of a possible future fleet of supersonic aircraft flying in the stratosphere on the
ozone layer could be slightly negative, but with a considerable uncertainty range
[10, 11].Increases in UV-B due to depletion of ozone in higher level may affect the
growth, photosynthesis and reproduction of phytoplankton. Reductions in
phytoplankton may result in a reduced uptake of CO2 in the oceans [12].
1.9 Radiative force- use of climate change prediction
Radiative forcing is a measure of how the energy balance of the Earth-
atmosphere system is influenced when factors that affect climate are altered. The
word radiative arises because these factors change the balance between incoming
solar radiation and outgoing infrared radiation within the Earth’s atmosphere. This
radiative balance controls the Earth’s surface temperature. The term forcing is used to
indicate that Earth’s radiative balance is being pushed away from its normal state.
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Radiative forcing is usually quantified as the ‘rate of energy change per unit area of
the globe as measured at the top of the atmosphere’, and is expressed in units of
‘Watts per square metre’. Ramaswamy[13] defines it as ‘the change in net (down
minus up) irradiance (solar plus long wave; in Wm-2
) at the tropopause after allowing
for stratospheric temperatures to readjust to radiative equilibrium, but with surface
and tropospheric temperatures and state held fixed at the unperturbed values’. When
radiative forcing from a factor or group of factors is evaluated as positive, the energy
of the Earth-atmosphere system will ultimately increase, leading to a warming of the
system. In contrast, for a negative radiative forcing, the energy will ultimately
decrease, leading to a cooling of the system.
Radiative forcing can be related through a linear relationship to the global
mean equilibrium temperature change at the surface (ΔTs):
ΔTs = λRF ( 1.1)
Where λ is the climate sensitivity parameter.
(1.2)
This equation, developed from these early climate studies, represents a linear view of
global mean climate change between two equilibrium climate states. Radiative
forcing is a simple measure for both quantifying and ranking the many different
influences on climate change; it provides a limited measure of climate change as it
does not attempt to represent the overall climate response.
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1.9.1 Radiative forcing for important green house gases
The important green house gases radiative forcing are shown in table 1.1
Table 1.1 Radiative forcing for green house
Trace gas Simplified Expression Radiative forcing,
ΔF(Wm-2
)
Constants
CO2
ΔF = α*+,-.-)/
ΔF= α*+,-.-)/0 1, /
ΔF = α,2,-/32,-)//
4567682,-/9*+,(0(: -0):))#- 0(:";()
3$-!/
α =5.35
α=4.841,
19):)')$
α=3.35
-<" ΔF= α, /3,=,>8?@A/ 3 =,>A ?@A// α =0.036
@ B ΔF= α, /3,=,>A ?@/ 3 =,>A ?@A// α =0.12
-C-3((!
ΔF = α(X-Xo) α =0.25
-C-3( ΔF = α (X-Xo) α =0.32
F(M,N)=0.47 ln[1+2.01X10-5
(MN)0.75
+5.31X10-15
M(MN)1.52
] (1.3)
Where C is CO2 in ppm, M is CH4 in ppb, N is N2O in ppb,X is CFC in ppb
The constant in the simplified expression for CO2 for the first row is based on
radiative transfer calculations with three-dimensional climatological meteorological
input data [14]. For the second and third rows, constants are derived with radiative
transfer calculations using one-dimensional global average meteorological input data
from [15,16] respectively. The subscript 0 denotes the unperturbed concentration. The
same expression is used for all CFCs and CFC replacements.
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Fig 1.5- Global mean radiative force values
For tropospheric ozone,
(ΔT2X)O = GO(ΔF2X) .. . . . (1.4)
Where GO = T/ (1-αp) So is the gain of the climate system with zero feedback
[16,17,18]. Taking the global-mean temperature T= 288K, planetary albedo αp= 0.3,
and solar irradiance So = 1367 Wm-2
yields GO = 0.3oC /Wm
-2.For ΔF2X =3.71 Wm
-2,
ΔT2X =1.12oC.
Total radiative forcing at time” t” [19], then
(1.5)
equation (1.5) assumes that the development of total radiative forcing is proportional
to radiative forcing due to carbon dioxide. Scaling parameter ω is a measure for the
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relative contribution of CO2 (1/ω is the share of CO2). Fig 1.5 shows global mean
radiative force values.
1.9.2 Direct and Indirect Radiative Forcing Agents:
Direct radiative forcing agents are emitted into the atmosphere and force the
climate system directly. Indirect radiative forcing agents act on the global
distributions of methane and tropospheric ozone and force the climate system
indirectly. Tropospheric ozone precursors such as VOCs, CO, H2, NOx are indirect
radiative forcing agents. Methane also acts as an indirect radiative forcing agent
although it is also a direct radiative forcing agent.
OH + VOC = RO2 (1.6)
RO2 + NO = NO2 + products (1.7)
NO2+ sunlight =NO +O3 (1.8)
NO+HO2=NO2+OH (1.9)
OH+CH4=products (1.10)
1.9.3 Gases: Hydroxyl Radical (OH): The detergent of the
atmosphere
OH is a major tropospheric oxidant. It removes CO/VOCs, is involved in
tropospheric ozone (O3) production, and in aerosol formation. It is the major sink of
CH4 in the atmosphere: OH determines CH4 lifetime. Multi-model OH highest in low
latitudes, especially over polluted regions. Changes in the future mostly negative, due
to large methane increases (sink) in this drastic scenario.
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1.9.4 Aerosols: major components
Sulphate (SO4) (anthropogenic and natural; natural comes mainly from
oceans and volcanoes). Black carbon (BC) (mostly anthropogenic; also from natural
fires). Organic carbon (anthropogenic and natural; natural comes from secondary
aerosol formation above forests). Mineral dust (mainly natural). Sea-salt (natural)
Nitrate (both anthropogenic and natural). Sulphate particles are produced from gases
(through OH oxidation) in the atmosphere. Their main precursors are:
a) anthropogenic or volcanic sulphur dioxide (SO2), b) dimethyl sulfide (DMS) from
biogenic sources, especially marine plankton. Sulphate is mostly scattering (cooling).
Black carbon is emitted in aerosol form (no gas precursors). It mainly comes from
fossil fuel combustion and biomass burning. BC is mostly absorbing (warming).
1.9.5 Drawback of Global radiative force in the case of aerosol and
surface ozone
Global radiative force forcing is not always useful, as: Temperature response
depends on a variety of uncertain feedbacks, and is highly region-dependent. Many
forcing agents, such as aerosols and tropospheric ozone (short-lived) are very
inhomogeneous, leading to complex patterns of forcing and response. A global view
of composition and radiation from satellites and from composition-climate models
(both recent developments!) can facilitate the study of such problems.
1.10 Stratospheric ozone and Temperature
Stratospheric ozone is not an anthropogenic forcing agent; rather it is (like
water vapour) an internal property of the atmosphere. For example, cooling of the
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upper stratosphere increases ozone abundance, by slowing ozone destruction rates,
and the increased ozone abundance offsets some of the cooling [20,21]. This ozone-
temperature feedback has, therefore, mitigated the upper stratospheric cooling that
would otherwise have occurred from the CO2 changes alone. The primary
anthropogenic forcing agents for past stratospheric ozone and temperature changes
are ozone-depleting substances (ODSs) and CO2 [22]. Because of the ozone-
temperature feedback, the upper stratospheric ozone depletion due to ODSs has been
partially masked by CO2-induced ozone increases. It follows that using the observed
ozone decreases to attribute global mean temperature changes must under estimate
the cooling due to ODSs (via ozone depletion), and overestimate the cooling due to
CO2 increases.
1.11 Stratospheric ozone and rainfall mechanism
Rainfall is an important parameter of our environment. O3 concentration is
maximum at a height of about 25 km above sea level. UV rays contain high-energy
photons which decompose ozone into oxygen. Ozone of the stratosphere absorbs
maximum of the incoming UV photons resulting a warm ozonosphere. But, it resists
the photons to fall directly on the clouds of the troposphere whose maximum height is
about 18 km above from the sea level. Hence it resists the unwanted warming of
cloud by UV rays and the high concentration of ozone in the stratosphere helps heavy
rainfall. The net reaction for the formation of ozone in the stratosphere is given
below:
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O2 + O2 → O3 + O (1.11)
O + O2 + M → O3 + M (1.12)
────────────────
3 O2 → 2 O3 (1.13)
It is shown by Midya [23] that if we consider that the reaction is endothermic,
this helps us to explain the dramatic decrease of ozone concentration over Antarctica
during spring time. So the higher temperature of the stratosphere will favor the higher
equilibrium concentration of ozone. Thus when the temperature of the stratosphere
increases by absorbing solar radiation, this will try to attain higher ozone
concentration of stratosphere. This also helps to attain low temperature of troposphere
and heavy rainfall. When stratospheric ozone is depleted, solar UV ray falls directly
on troposphere. As a result temperature of stratosphere decreases and that of
troposphere increases. This helps to attain an unfavorable condition of rainfall.
Although this is the main mechanism of rain fall, less intense rain (showers) can
occur even when the cloud does not contain ice crystals at all. But, the presence of
CCN is essential for heavy rainfall. One study says that, stratospheric ozone
concentration must be at a certain concentration level in order to obtain sufficient
rainfall. One study says that ozone concentration decreases, monsoon rainfall
increases with time, and monsoon rainfall occurs only when ozone concentration
reaches to a certain concentration level along with other parameters of precipitation.
This level slightly differs from station to station and it confirms latitudinal variations
of ozone concentrations for which rainfall takes place [24]. Monsoon rainfall is a very
complex process. It not only depends on ozone concentration but also on CCN,
humidity, temperature, El-Nino, etc. Monsoon rainfall starts when ozone
concentration reaches to a certain concentration level and monsoon rainfall decreases
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with the increase of ozone concentration. Midya [25] showed that N2O plays
important role in the depletion of stratospheric O3. Decrease of O3 concentrations
over India may be due to the increase of N2O over India.
1.12 Stratospheric ozone and UV radiation
Sunlight consists of solar rays of differing wavelengths. Visible light ranges
from 400nm (violet) to 700nm (red). Infrared radiation, or heat, has longer
wavelengths than visible light; ultraviolet radiation has shorter wavelengths than
visible light. UVR is further divided into UVA (315–400nm), UVB (280–315nm) and
UV-C (<280nm). Almost all incoming solar UVC and 90% of UVB are absorbed by
stratospheric ozone, while most UVA passes through the atmosphere unchanged.
Although UVA penetrates human skin more deeply than UVB, the action spectra
from biological responses indicate that it is radiation in the UVB range that is
absorbed by DNA—subsequent damage to DNA appears to be a key factor in the
initiation of the carcinogenic process in skin. The amount of ambient UVB
experienced by an individual outdoors with skin exposed directly to the sky is
dependent on the following:(i) stratospheric ozone levels(ii) solar elevation(iii)
regional pollution(iv) altitude of the individual(v) cloud cover(vi) presence of
reflective environmental surfaces such as water, sand or snow. Since ozone is not the
only influence on surface UV, and may not be the only parameter to change in the
future, or to have done so in recent years. Clouds can cause large and rapid changes
(increases as well as decreases) in surface UV radiation and in total radiation (i.e.,
radiation integrated over the entire spectrum from about 300 nm to 3000 nm). On
average, clouds have an attenuating effect of 15–32% in the UV wave band [26, 27],
but for any given situation the cloud effect will depend on the cloud type, depth, and
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distribution across the sky. Thus, cloud transmission of UV irradiance is difficult to
quantify in sufficient detail to provide an exact determination of its effect at a given
time and place [28].
Solar uv rays contain high energy photons which decompose water molecules
into atomic oxygen and OH radical. Thus we can expect that it stratospheric ozone is
depleted, the high energy solar uv photon fall directly on troposphere and as a result,
tropospheric water molecules are decomposed resulting the fall of relative humidity
and increase of tropospheric ozone which acts as greenhouse gas and plays important
role in global warming [29].
1.13 Stratospheric ozone and EL NINO
In terms of disturbance, the area is subject to hurricanes and tropical storms
and the increased precipitation and winds from these events. Disturbances most often
arise from (i) cyclonic systems, (ii) noncyclonic inter tropical systems, (iii) extra
tropical frontal systems, and (iv) large- scale, coupled ocean-atmospheric events (e.g.,
North Atlantic Oscillation, El Niño-Southern Oscillation)[30]. Intratropical
atmospheric systems start and stay in the tropics. For large-scale ocean-atmospheric
systems such as the NAO and ENSO, the NAO has the stronger relationship with
PR’s annual climate than ENSO. A high winter NAO is correlated with lower
precipitation in PR. ENSO events are correlated with temperature in the Luquillo
Mountains and associated with wetter conditions from April to July. El Nino is
associated with drier conditions from September to October throughout the
Caribbean. All of these relationships are generally fairly weak, however, and have a
poor link to hurricanes and other large-scale disturbances.
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1.14 Stratospheric ozone and Quasi Biennial oscillation
Stratospheric ozone is produced by ultraviolet photo dissociation of molecular
oxygen and destroyed by catalytic reactions of the oxides of hydrogen, nitrogen,
chlorine, and bromine. The distribution of ozone with latitude, longitude, and altitude
is determined by large and small-scale transport between regions of net production
and regions of net loss. Stratospheric ozone is sensitive to variability in both
photochemical and dynamical processes and thus varies on seasonal and longer time
scales. After accounting for variability due to the seasonal and solar cycle in
ultraviolet flux, variability due to volcanic aerosols, and variability due to dynamic
influence such as the quasi-biennial oscillation(QBO), analysis of observations
indicate a downward trend in stratospheric ozone over the time period from1979 to
the late 1990s [31,32,33]. The growth of reactive chlorine and bromine in the
stratosphere due to decomposition of industrially produced chlorofluorocarbons
(CFCs), halons, and methyl bromide has, over the last two decades, led to a decrease
in the amount of ozone overhead. The statistical analysis accounts for random and
coherent dynamical variability other than the QBO only in the noise term that helps to
evaluate uncertainties.
1.15 The Formation of Tropospheric Ozone
Tropospheric ozone, which is the most important oxidant in the atmosphere, is
formed through in-situ photochemical reactions involving nitrogen oxides and
volatile organic compounds [34-37]. Enhanced levels of tropospheric ozone are a
serious concern because of effects on human health, [38-42] crops, ecosystems,
building materials [43-45] and, hence, the economy [46-51]. Additionally,
tropospheric ozone is the third most important greenhouse gas after carbon dioxide
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and methane [52-53], hence, the increasing trend in tropospheric ozone is influencing
both the chemical composition of the atmosphere and global climate. The main
sources of ozone in the troposphere are: Incursion from the stratosphere. Reactions
involving precursors that come from biogenic volatile organic compounds (VOC);
and, Photochemical reactions with the precursors, nitrogen oxides (NOx = NO +
NO2), carbon monoxide (CO), methane (CH4) and other organic compounds resulting
from anthropogenic activities. It was thought that tropospheric ozone largely resulted
from the processes of stratosphere-troposphere exchange and organic materials
deposition at the Earth’s surface.
As far as the photochemistry is concerned, the formation mechanism is
influenced by the mixing ratio between all the relevant chemical compounds (for
example, precursors and oxidants). Otherwise, an important factor related to ozone
formation is solar radiation with a wavelength range of around 320nm to 410nm [54].
A general process of ozone formation at the boundary layer can be written as follows
[55]:
NOx + NMHCs + hν+ M (N2, O2) → O3 + other photochemical oxidants (1.14)
(NMHCs = non methane hydrocarbons, hν= ultraviolet solar radiation,
320nm~ 410nm, and other terms have been defined above).
A conversion reaction of NOx should be noted, involving a catalytic cycle of
ozone formation and consumption. The cyclic reactions are as follow:
NO2 + hν→ O + NO (1.15)
O + O2 + M → O3 + M (M = N2, O2) (1.16)
O3 + NO → O2 + NO2 (1.17)
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The major physical mechanisms affecting tropospheric ozone formation include local
meteorological conditions (such as humidity, temperature, wind speed and direction)
[56, 57]. For instance, low humidity, low wind speeds, calm weather and cloud-free
conditions favour the photochemical reactions. Therefore, the precursors and ozone
from the vicinity could bring more serious effects on the air quality. The lifetime of
ozone in the troposphere is typically a few weeks [58].
1.15.1 National Ambient Air Quality Standards for ozone
Ambient standards and guidelines for ground-level ozone or surface ozone are
aimed at protecting human health, sensitivity ecosystems, and agricultural plants from
the harmful effects of ground-level Ozone. Final rules signed June 2010. National
Ambient Air quality standard had attained by using three year average of the daily
maximum one hour average value and each value must not exceed 75 ppb
(0.075ppm).Table 1.2 gives the prescribed values of surface level ozone as per
National Ambient Air Quality Standards.
Table 1.2. National Ambient Air Quality Standards (NAAQS)
Pollutant Level Averaging Time
Ozone
0.075 ppm
(2008 std) 8-hour
0.08 ppm
(1997 std) 8-hour
0.12 ppm 1-hour
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1.16 Tropospheric ozone and Clouds
Clouds are condensed atmospheric moisture in the form of minute water
droplets or ice crystals. Formation and destruction of ozone in the troposphere
comprises a series of complex cycles in which atomic oxygen, molecular oxygen,
carbon monoxide, oxides of nitrogen, water vapour, volatile organic compounds, etc.
are involved. Ozone in the troposphere is produced by the addition of ground state
oxygen atoms O (3p) to molecular oxygen assisted by any third body M to ensure
simultaneous momentum and energy conservation [59].
O3p_+O2 +M → O3 +M. (1.18)
The primitive terrestrial atmosphere was oxygenic in nature and ozone was produced
by photo dissociation of water vapour as follows [60].
H2O+hν → OH + H (1.19)
OH + hν → O+H (1.20)
O+O+M → O2+M (1.21)
O+O2 +M → O3 +M (1.22)
1.17 Tropospheric ozone and temperature, Humidity
Temperature, atmospheric humidity and sunshine levels influence background
O3 levels through the controls they exert on the photochemistry of net O3 production.
Temperature is often used as a predictor for high O3 [61, 62]because of its direct
influence on chemical kinetic rates and the mechanism pathway for the generation of
O3 [e.g., H-abstraction versus OH addition [63] and strong correlation with stagnant,
sunny atmospheric conditions .Higher temperatures accelerate O3 production
(depending on the NOX regime), and especially in summer, are also likely to increase
biogenic VOC emissions and hence lead to higher surface O3 concentrations in high
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NOX regions. Temperatures are expected to increase more over the continents than
the oceans. The chemistry of O3 production, initiated by reactions of OH with CH4,
CO and VOC. OH is produced following photolysis of O3 in the presence of water
vapour, therefore absolute humidity is a strong determinant of background O3 levels,
and the expected changes in atmospheric humidity associated with global warming
are likely to have complex effects on surface O3.
1.18 Tropospheric Ozone and ENSO
The dramatic variation of seasonal distributions of high ozone days seems to
be related to short-term variation of climate that might result from the effect of El
Niño Southern Oscillation (ENSO). McPhaden [64] indicated that the year 1997-1998
was the strongest El Niño phenomenon on record and followed by an extended La
Niña period that began in mid-1998 to the winter of 2000 and slowly decayed from
the spring of 2001. The interplay between the Asian winter monsoon and ENSO plays
an influential role in East Asian climate. The linkage between the climate of East
Asia and ENSO has been found, for example, a weaker East Asian winter monsoon
along the East Asian coast occurs during the mature phase of ENSO and the
relationship between the rainfall patterns and ENSO. Nevertheless, the effects of La
Niña on climate show a reverse situation, especially in equatorial region [65].
1.19 Tropospheric ozone and Lightning
Roughly 10% of the ozone in the troposphere plays a key role in providing
oxidizing radicals, OH that cleanses the lower atmosphere of pollutants, such as
carbon monoxide (CO), and sulfur dioxide (SO2). Global NOx (NO+NO2) has various
sources, including fossil fuel combustion, biomass burning, aircraft emission,
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microbial processes in the soil, and lightning .Nitric oxide is produced in the hot
lightning channel by way of the Zel’dovich mechanism, which considers the breakup
of N2 and O2 molecules [66]. This mechanism takes place due to the high
temperatures and pressures initiated by the lightning stroke. Lightning, which is
calculated internally in the model, contributes about 16% to the NOx global budget;
while fossil fuel is burning and biomass burning contributes close to 70% of the NOx
budget. Measured concentrations of NOx have substantial spatial and temporal
variability. This is due to the fact that the local chemical lifetime of NOx is less than
1 day in the polluted boundary layer and 5 to 10 days in the upper troposphere. The
effect of lightning on the NOx budget is important to NOx in the upper troposphere,
where the lifetime is longer. Many studies have looked at lightning produced NOx
[67].
1.20 Tropospheric Ozone is toxic to plants and human healths
Recent studies show that ozone is a big problem around the world which
affects vegetation and humans and contribute also to climate change. Tropospheric
ozone is the major ingredient in photochemical smog and, as mentioned earlier,
represents a considerable risk to vegetations and human beings. Effects of ozone may
occur at various levels of organization, i.e. from the cellular level through the level of
individual organs and plants to the level of plant communities and ecosystems [68].
Fig 1.6 shows climate change direct and indirect effects of ozone depletion.
Growth reduction from chronic exposures as well as crop yield losses,
reductions in annual biomass increments for forest trees and shifts in species
composition of semi-natural vegetation are other well-documented effects of ground-
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level ozone or surface ozone [69]. Current levels of ozone reduce yields of major
staple crops such as rice, wheat, corn, potato and soybean.
Fig1.6 Health impacts due to climatic changes and ozone layer depletion
1.21 Tropospheric ozone driver of climate change
Ground-level O3 increases also contribute to climate change through both
direct and indirect routes. Ozone concentrations should therefore be considered
alongside those of rising CO2 concentrations. The IPCC estimates that tropospheric
O3 increases since pre-industrial times have contributed somewhere between 0.25 and
0.65 W m–2 to global radiative forcing [70]. Increases in O3 will also have an indirect
effect on global warming by suppressing plant growth, reducing the land carbon sink
for CO2 and therefore increasing the rate at which CO2 increases in the atmosphere.
Experimental studies have shown that O3 reduces the additional carbon storage
arising from increased CO2 concentrations, but that elevated CO2 concentrations can
reduce the negative impacts of O3 on vegetation. A physiological model linking these
phenomena has recently demonstrated that the indirect radiative effects of O3 via
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reduced carbon sequestration could increase the total radiative forcing due to O3 over
the period 1900–2100 by at least 70% [71]. This suggests that tropospheric O3
increases are an even more important driver of global warming.
This suggest that future changes in climate may not have a major influence on
O3 concentrations globally, but that it is more important at the regional and local
scales. However, many of the climate processes which control O3 at these scales are
not yet captured well within global atmospheric chemistry models. This provides a
brief evaluation of which climatic processes may influence O3 concentrations in the
future. It also highlights the role of tropospheric O3 as a greenhouse gas. The
production of O3 is controlled by temperature, sunlight and humidity, and by the
long-range transport of pollutants all of which are sensitive to changes in climate.
1.22 Ozone effects on other things
There are further environmental effects of ozone, for example: Rubber, textile
dyes, fibers, and certain paints may be weakened or damaged by exposure to ozone.
Some elastic materials can become brittle and crack, while paints and fabric dyes may
fade more quickly. It also damages cotton, acetate, nylon, polyester, and other
textiles. Reactions involving ozone also cause deterioration of electronic devices.
1.23 Tropospheric ozone forcing future trend
1850-2000 forcing is mostly positive, except for the Antarctic. It peaks in the
northern subtropics. 2000-2100 tropospheric ozone forcing is large in the scenario
with large methane changes.
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1.24 Relation between stratospheric ozone and tropospheric ozone
One recent studies says that the nature of long term seasonal and yearly
variations of tropospheric and stratospheric ozone over thumba (8.5oN,77
oE) and
bangalore (13oN,77.5
oE) ,India. Seasonal variations clearly reveal that stratospheric
ozone concentration attained comparatively higher value in the months from july to
September while lower tropospheric ozone concentration at the same time over above
stations during period 1979-2005.Yearly variations shows increasing trend in
tropospheric ozone but decreasing trend in stratospheric ozone from 1979 to 2005.
Undesirable environmental effects due to such tropospheric rise and stratospheric
decline in ozone are also mentioned [72].
1.25 Total Ozone Column (TOC)
The most common ozone measurement unit is the Dobson Unit (DU). The
Dobson Unit is named after atmospheric ozone pioneer G.M.B. Dobson who carried
out the earliest studies on ozone in the atmosphere from the 1920s to the 1970s.
Fig 1.7 shows the TOC measured by Dobson unit. A DU measures the total amount
of ozone in an overhead column of the atmosphere.
Dobson Units are measured by how thick the layer of ozone would be if it
were compressed into one layer at 0 degrees Celsius and with a pressure of one
atmosphere above it. Every 0.01 millimeter thickness of the layer is equal to one
Dobson Unit.
The average amount of ozone in the stratosphere across the globe is about
300DU (or a thickness of only 3mm at 0oC and 1 atmospheric pressure!). Highest
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levels of ozone are usually found in the mid to high latitudes, in Canada and Siberia
(360DU).
Fig 1.7 Total Ozone column measured by Dobson Unit
1.26 Objectives of this study
1) To study and analyze the role of surface level ozone or ground level ozone and
total ozone column activity in the climate changes.
2) To measure the surface ozone at different sites over Tamil Nadu. Based on the
measured data forecasting future ozone value.
3) Before forecast, need to do Rescaled Analysis and Chaotic analysis of ozone
data.
4) Depends upon the ozone analysis we need to develop linear and non-linear
mathematical models for estimation of surface ozone and total column ozone.
5) To correlate total column ozone and rainfall of Kodaikanal using the classical
decomposition method and Auto Regressive Integrated Moving Average
method.
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6) To develop a Artificial Neural Network and Adaptive Neuro Fuzzy Inference
System (ANFIS) model for estimation of surface ozone as output and wind
speed, temperature, relative humidity as input.
1.27 Thesis outline
This Thesis consists of six chapters.
The First Chapter deals with introduction and explains the formation of
ozone, national ambient air quality standards of ozone, various ozone units, and
different levels of ozone measuring technique, effects of ozone on climate change,
effects of ozone on vegetation, effects of ozone on human health and environmental
problems. Future ozone trend, Total ozone column are discussed in this chapter.
The Second Chapter describes briefly the study area and the data used in the
study. Surface ozone along with different meteorological parameters such as
temperature, relative humidity, and wind speed are also measured and taken for the
study. In the study site area Chennai and Kodaikanal were considered. In Chennai site
diurnal and seasonal variations were observed. In Kodaikanal total column ozone
along with rainfall data were gathered and analyzed.
The Third Chapter is analysis of Hurst, Fractal dimension and chaotic nature
of ozone data .Hurst (H) and fractal dimension (D) determines whether the given time
series is completely random or has some long term memory. Chaotic analysis is used
to find the power of predictability. Whether the time series can be predicted for short
term or long term is identified by the chaotic analysis.
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The Fourth Chapter enlightens ozone forecasting. Based on the analysis of
ozone data various forecasting methods used. In this chapter SDPV Forecasting
Method is used. SDPV consists of Classical Decomposition (CD), Auto Regressive
Integrated Moving Average (ARIMA) methods in ozone Forecasting.
The Fifth Chapter elucidates DIMP Forecasting Method. DIMP consists of
Artificial Neural Network (ANN) and Adaptive Neuro Fuzzy Inference System
(ANFIS) Forecasting model. Chennai surface ozone data was analyzed by ANN and
ANFIS model.
The Sixth Chapter summarizes the significant findings of the research work.
Some essential conclusions are drawn in this chapter. The performance of four
forecasting methods is found. The recommendations for further research are given in
this chapter. The relevant references in connection with this work are given at the end
of each chapter. In addition the lists of publications are presented in the appendix
section.
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