Rocket-borne in-situ
measurements in the
middle atmosphere
Jonas Hedin
AKADEMISK AVHANDLING
för filosofie doktorsexamen vid Stockholms Universitet
att framläggas för offentlig granskning
den 20:e februari 2009
Rocket-borne in-situ measurements in the middle atmosphere
Doctoral thesis
Jonas Hedin
Cover photo: Launch of the HotPay-I sounding rocket at Andøya Rocket Range, Norway
on July 1, 2006 (photo by Jonas Hedin).
ISBN 978-91-7155-813-8, pp. 1 - 59
© Jonas Hedin, Stockholm 2009
Stockholm University
Department of Meteorology
10691 Stockholm
Sweden
Printed by Universitetsservice US-AB
Stockholm 2009
2
Contents
Abstract
5
List of papers
7
1 Introduction
9
1.1 The Earth’s atmosphere
9
1.2 Important mesospheric phenomena and their study
12
1.2.1 Ice particles
13
1.2.2 Meteoric smoke
17
1.2.3 Nightglow
23
2 Atmospheric science with sounding rockets
26
2.1 The need for sounding rockets
26
2.2 Challenges
27
3 The PHOCUS project
30
3.1 Science questions
31
3.2 Instrumentation
35
4 Results of this thesis
39
5 Outlook
43
Acknowledgements
45
Bibliography
47
3
4
Abstract
The Earth's mesosphere and lower thermosphere in the altitude range 50130 km is a fascinating part of our atmosphere. Complex interactions
between radiative, dynamical, microphysical and chemical processes give
rise to several prominent phenomena, many of those centred around the
mesopause region (80-100 km). These phenomena include noctilucent
clouds, polar mesosphere summer echoes, the ablation and transformation
of meteoric material, and the Earth’s airglow. Strong stratification and
small scale interactions are common features of both these phenomena and
the mesopause region in general. In order to study interactions on the
relevant spatial scales, in situ measurements from sounding rockets are
essential for mesospheric research.
This thesis presents new measurement techniques and analysis methods for
sounding rockets, thus helping to improve our understanding of this remote
part of the atmosphere. Considering the need to perform measurements at
typical rocket speeds of 1 km/s, particular challenges arise both from the
design of selective, sensitive, well-calibrated instruments and from
perturbations due to aerodynamic influences. This thesis includes a
quantitative aerodynamic analysis of impact and sampling techniques for
meteoric particles, revealing a distinct size discrimination due to the
particle flow. Optical techniques are investigated for mesospheric ice
particle populations, resulting in instrument concepts for accessing smaller
particles based on Mie scattering at short ultraviolet wavelengths. Rocketborne resonance fluorescence measurements of atomic oxygen are critically
re-assessed, leading to new calibration concepts based on photometry of O2
airglow emissions.
The work presented here also provides important pre-studies for the
upcoming PHOCUS rocket campaign from Esrange in July 2010. PHOCUS
(Particles, Hydrogen and Oxygen Chemistry in the Upper Summer
mesosphere) will address the interaction between three major mesospheric
players: meteoric smoke, noctilucent clouds and gas-phase chemistry.
5
6
List of papers
This thesis consists of an introduction and the following five papers
I
Hedin, J., Gumbel, J., and Rapp, M.: On the efficiency of rocket-borne
particle detection in the mesosphere, Atmospheric Chemistry and
Physics, 7, 3701-3711, 2007.
II
Hedin, J., Gumbel, J., Waldemarsson, T., and Giovane, F.: The
aerodynamics of the MAGIC meteoric smoke sampler, Advances in
Space Research, 40, 818-824, 2007.
III Rapp, M., Hedin, J., Strelnikova, I., Friedrich, M., Gumbel, J., and
Lübken, F.-J.: Observations of positively charged nanoparticles in the
nighttime polar mesosphere, Geophysical Research Letters, 32,
L23821, doi:10.1029/2005GL024676, 2005.
IV Hedin, J., Gumbel, J., Khaplanov, M., Witt, G., and Stegman, J.:
Optical studies of noctilucent clouds in the extreme ultraviolet,
Annales Geophysicae, 26, 1109-1119, 2008.
V
Hedin, J., Gumbel, J., Stegman, J., and Witt, G.: Use of O2 airglow for
calibrating direct atomic oxygen measurements from sounding
rockets, manuscript for submission to Atmospheric Measurement
Techniques, 2009.
7
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Chapter 1
Introduction
1.1 The Earth’s atmosphere
The Earth’s atmosphere (Greek ‘atmos’ = ‘vapour’ + ‘sphaira’ = ‘ball’) can
be divided into regions and boundaries, or “spheres” and “pauses”, in
several ways based on the properties the atmosphere is described in
[Brasseur and Solomon, 1986]. The most common way is a division
according to the temperature change with altitude (figure 1). The lowermost
region called the troposphere (Greek ‘tropos’ = ‘turning’ or ‘change’) is
where our daily weather takes place. It extends to about 7 km at the poles
and about 17 km at the equator and is characterised by rather strong vertical
mixing, hence its name. Molecules can traverse the entire depth in periods
between a few days and a few minutes depending on the weather situation.
The mixing is due to solar heating of the surface that in turn heats the air
masses which then become less dense and rise. While rising, it cools
adiabatically and latent heat can be released as water vapour condenses,
leading to a vertical temperature decreases of typically ~ 6.5 K/km up to the
tropopause, where a local temperature minimum of about 200 - 230 K is
reached depending on latitude. In the stratosphere (Latin ‘stratus’ =
‘layered’) the temperature increases with altitude due to the absorption of
shortwave solar ultraviolet (UV) radiation by ozone (O3) (λ ~ 200 - 300
nm). The main cooling agent in the stratosphere is the ν2 band (bending
vibrational state) of carbon dioxide (CO2) at infrared (IR) wavelengths of
15 µm. The stratosphere extends up to the stratopause at about 50 km and
is characterised by very small vertical mixing with transport time-scales of
the order of years. The stratosphere is very dry since only about 3 ppmv of
water vapour is transported from the troposphere past the ‘cold trap’ at the
tropopause. An additional ~3 ppmv of water vapour is produced in the
9
stratosphere from the oxidation of methane (CH4), transported from the
Earth’s surface.
120
Thermosphere
Heterosphere
turbopause
100
mesopause
Altitude [km]
80
Mesosphere
Summer profile
Vinter profile
60
stratopause
40
20
Homosphere
Stratosphere
tropopause
Troposphere
0
100
150
200
250
Temperature [K]
300
350
Figure 1. The temperature structure of the atmosphere.
In the mesosphere (Greek ‘mesos’ = ‘middle’) ozone levels and, hence,
direct shortwave heating at 200 – 300 nm is low. As a consequence,
adiabatic temperature decrease with altitude is dominant again, as in the
troposphere. In the upper mesosphere, between about 80 and 100 km, the
mesopause region forms a boundary to the thermosphere above. In summer
this boundary is more sharply defined, with temperature minimum that at
polar latitudes can reach extremely low temperatures (down to 100 K)
around 90 km [Höffner and Lübken, 2007], while during other seasons the
minimum (~180 - 200 K) is less well defined and can occur over a range of
several 10 km. Above the mesopause, in the thermosphere (Greek
‘thermos’ = ‘heat’), the absorption of extreme ultraviolet (EUV) and X-ray
radiation at wavelengths below 180 nm, mostly by O2, leads to a rapid
warming with altitude.
The summertime polar mesopause is the coldest place on Earth and it is
here that ice particles form and can be seen optically as noctilucent clouds
in twilight conditions. It is also in this region where meteoroids deposit
much of their material as they burn up entering the atmosphere, and excited
atoms and molecules emit radiation known as nightglow (or more general
airglow) (see chapter 1.2.3 and paper V). The reason why the upper
10
mesosphere is so cold in summer and warmer in winter is the wave-driven
general circulation. A detailed review of gravity wave dynamics and its
effects in the middle atmosphere has been given by Fritts and Alexander
[2003] Gravity waves generated in the lower atmosphere (e.g. in flow over
topography, convection, wind shear, geostrophic adjustment and wavewave interactions) propagate upwards and eventually become unstable and
break, depositing momentum and energy in the mesosphere. This wave
forcing drives the atmospheric flow velocity into the direction of the wave
phase velocity. In summer this acts as a “wave drag” on the prevailing
zonal winds, slowing down the mean wind near the mesopause and
completely overwhelms the radiative forcing which otherwise would
maintain very high easterly wind speeds at mesopause heights. The wave
drag thus forces a reversal of the zonal winds from easterly at lower heights
to westerly in the lower thermosphere. It also forces a mean equatorward
wind of the order of 10 m/s in the summer and a similar poleward flow in
the winter hemisphere. The equatorward moving air, in combination with
the radiative forcing, induces strong upward motion (of the order of several
cm per second) over the summer pole as a result of mass continuity. The
rising air cools adiabatically, resulting in the very low temperatures of the
summer mesopause while the opposite circulation occurs in the other
hemisphere, and downwelling warms the wintertime mesosphere.
Another important way of describing the atmosphere is to divide it
according to how the atmospheric composition changes with altitude.
Composition changes occur with altitude both because of transport
properties and because of chemical properties. As for transport, the altitude
range from the Earth’s surface up to the turbopause at about 100 km is
called the homosphere as it features more or less uniform conditions due to
turbulent mixing of the atmosphere. However, above the turbopause, in the
heterosphere, the composition starts to vary with altitude. This is because
in the absence of turbulent mixing the rate of fall-off in density of a species
with increasing altitude depends on the molecular mass. Hence, heavier
atmospheric constituents such as molecular oxygen and nitrogen fall off
more quickly than lighter constituents such as atomic oxygen, hydrogen
and helium. The absence of turbulence above ~100 km is closely related to
the increasing molecular mean free path with altitude. This increase of the
mean free path does not only have a major effect on atmospheric
composition but also on aerodynamics, as will be discussed in papers I and
II. As for chemistry, composition changes with altitude as sources and sinks
for various species are highly variable. In the middle and upper atmosphere,
this is in particular true for the flux of solar UV radiation that provides
much of the energy for atmospheric photochemistry. This is the basis e.g.
for the production of odd oxygen, the partition between atomic oxygen and
11
ozone, and the generation of excited airglow species that will be discussed
in paper V.
The ionosphere is the altitude region of the atmosphere where ionisation of
atoms and molecules by shortwave solar UV radiation occurs. It includes
both the mesosphere and the thermosphere. The ionosphere can be divided
up into several layers depending on the main ionised species. In the
mesosphere, NO is ionised by the very intense solar Lyman-α line at 121.57
nm, giving rise to the ionospheric D layer. Photolysis by Lyman-α is also
the main loss process of water vapour. The E-layer between 90 and 120 km
is formed by soft X-ray and far UV ionisation of O2 and the F-layer above
120 km by EUV ionisation of O. The interaction of ionospheric charges
with atmospheric particles will be discussed in paper III.
Figure 2. Schematic overview of the middle atmosphere processes (water vapour
in ppm(v)).
1.2 Important mesospheric phenomena and their
study
The mesosphere and lower thermosphere (MLT) region of the Earths
atmosphere between 50 and 130 km is a transition region that is controlled
by a combination of radiative and dynamic forcings from above and below
(figure 2). Prominent phenomena related to the complex interactions in this
region are noctilucent clouds (NLC) [Rapp and Thomas, 2006], polar
12
mesosphere summer echoes (PMSE) [Rapp and Lübken, 2004], the ablation
and transformation of meteoric material [Plane, 2003] and the Earth’s
nightglow [Meriwether, 1989]. Systematic observations of NLC, PMSE,
meteoric smoke or the nightglow provide important clues to our
understanding of the middle atmosphere as a whole. However, in order to
draw robust conclusions from such observations, we need a more robust
understanding of the physical and chemical processes behind these
phenomena.
Figure 3. Noctilucent clouds observed over Stockholm. (Photo by Nathan Wilhelm)
1.2.1 Ice particles
NLC’s (figure 3) and PMSE’s are related to ice particles in the polar
summer mesopause region. Existing just at the edge of feasibility,
mesospheric ice clouds are expected to be extremely sensitive to changes in
middle atmospheric conditions and have thus gained interest as indicators
of global variability and trends. It has been argued that even small longterm changes of mesospheric water vapour (e.g. due to anthropogenic
methane emissions or changes in lower atmospheric circulation patterns) or
mesospheric temperatures (e.g. due to anthropogenic carbon dioxide
emissions) should lead to prominent long-term changes of the observed
13
properties of mesospheric ice clouds [Thomas et al., 1989]. The question
whether such long term changes are already evident in the experimental
records has been debated [von Zahn, 2003; Thomas et al., 2003]. In
addition, the occurrence of mesospheric ice particles has been discussed in
connection with space traffic exhaust [Stevens et al., 2003, 2005], observed
differences between the Arctic and Antarctic mesosphere [Hervig and
Siskind, 2005], and the middle atmosphere coupling between the two
hemispheres [Becker et al., 2004; Karlsson et al., 2007].
We know today that mesospheric clouds are composed of water ice
particles, possibly with small nucleation cores. A detailed review about
current knowledge of the microphysics of mesospheric ice particles has
been given by Rapp and Thomas [2006]. Ice particle nucleation takes place
at the altitudes with the largest saturation ratios S=pH2O/pSat (where pH2O is
the water vapor partial pressure and pSat is the equilibrium vapor pressure of
water vapor over ice). This is the case is at altitudes close to the
temperature minimum between ~86 and 90 km. Lübken [1999] found that
typical saturation ratios in the mesopause region are on the order of ~100
(although significantly higher supersaturations can be expected for wavedisturbed conditions). Thus, homogeneous nucleation should be negligible,
and pre-existing ice nuclei are required for the formation of ice particles.
Once formed, the ice particles then continue to grow by direct deposition of
water vapor onto their surface. They settle down due to gravity and
consume the available water vapor on their way down through the
atmosphere. During this motion, the particles are further subjected to
transport by mean winds and small scale motions, i.e. waves and
turbulence. Once the particles have reached equivalent spherical radii in
excess of ~30 nm, they scatter light efficiently such that they may
eventually be observed optically with the naked eye or from ground based
or space borne optical instruments as NLC or polar mesospheric clouds
(PMC).
All these ice particles are also immersed in the plasma of the ionospheric
D-region. Hence, electrons attach to the ice surfaces so that the particles
become charged. In addition, the 80 – 90 km altitude range is the region in
the atmosphere where gravity waves propagating from below grow unstable
and produce turbulence. The charged ice particles are transported by the
turbulent velocity field, leading to persistent small scale structures in the
ionospheric charge distribution. This causes irregularities in the radio
refractive index (which is effectively determined by the electron number
density at these altitudes) which can be observed from the ground by radars
as PMSE [Rapp and Lübken, 2004]. Important is that the lifetime of the
irregularities against diffusion is determined by the radius of the ice
14
particles. Typical values for the lifetime are just a few minutes for small
radii at altitudes above ~85 km and up to hours for the larger radii below.
As a consequence, PMSE at altitudes above typically 85 km can only exist
during phases of active neutral air turbulence. In contrast, the long lifetime
below this altitude “freezes” structures in the plasma structures so that
PMSE also can occur long after the disappearance of neutral air turbulence.
One central open question is the actual origin of the ice particles.
Homogeneous nucleation is very unlikely to occur at typical mesospheric
supersaturation levels, so there has to be some kind of nucleus that the
water molecules can nucleate on. Candidates that have been proposed in the
past are: large proton hydrate ion clusters [Witt, 1969], soot [Pueschel et
al., 2000], sulfuric acid aerosol particles [Mills et al., 2005], sodium
bicarbonate [Plane, 2000], sodium hydroxide [Petrie, 2004] and meteoric
smoke particles [Rosinski and Snow, 1961; Hunten et al., 1980]. The
meteoric smoke particles have long been considered as the most likely
candidate but there are uncertainties: The influx, or the amount of meteoric
material that enters the Earth’s atmosphere per year, has an uncertainty of
one to two orders of magnitude. The processes of meteoric ablation and recondensation into smoke particles are poorly understood. And, last but not
least, it has recently been recognized that the global circulation efficiently
transports meteoric material away from the polar summer mesopause
region [Megner at al., 2008a]
The initial step of ice particle formation, the nucleation process, introduces
the most significant uncertainty into the current theoretical understanding
of mesospheric ice microphysics. The nucleation feasibility and rate of ice
particle formation is linked to the type of ice nuclei, including composition
and corresponding properties like their crystal structure, their number
density and size distribution, as well as their charge state. The actual shape
of the particles is uncertain and this affects the growth rate and
sedimentation speed. In most modeling efforts the particles are considered
to be spherical, but as they are most probably non-spherical [Baumgarten et
al., 2002; Eremenko et al., 2005], the sedimentation speed will decrease
and hence the particle will spend more time in the altitude range with
favorable conditions for growth. Also the equilibrium vapor pressure of
water vapor over ice, or the saturation vapor pressure, is uncertain for the
temperatures relevant to the polar summer mesopause region. In addition,
the very nucleation process on nanometer-size nuclei is a major source of
uncertainty since the commonly applied liquid drop theory is of limited
validity under these conditions [Kessee, 1989; Gumbel and Megner, 2008].
Finally, the accommodation coefficient, i.e. the efficiency of collisional
energy transfer from ice particles and the surrounding air molecules, is
15
another major uncertainty that affects both ice particle temperature and
sedimentation speed.
The size range of visible NLC particles has been studied by optical
measurements with satellites [e.g. Rusch et al., 1991; Merkel et al., 2003;
Karlsson and Rapp, 2006; Bailey et al., 2008], sounding rockets [Gumbel et
al., 2001], and lidars [von Cossart et al., 1999; Baumgarten et al., 2007].
The size retrievals are based on the comparison of measured scattering
properties for sunlight (spectral dependence or phase function) to
theoretical simulations of the size dependence of Mie scattering. According
to these measurements the particle radii generally remain smaller than 100
nm. However, from normal optical measurements it is not possible to
access the smaller particles in the mesospheric ice population. This has two
reasons: (1) The scattering is strongly dominated by the larger particles in
the population because of the strong size dependence of the scattering. (2)
Smaller sizes become indistinguishable because they all fall into the
Rayleigh scattering regime, i.e. the particle radius is much smaller than the
light’s wavelength. The current lack of knowledge on the smaller particles
is a major limitation when studying important questions regarding the
nucleation and growth processes governing the ice particles in the
mesosphere.
Optical measurements are today the only means of obtaining information
about the size of NLC particles. New information about the properties of
the particles has become available from measurements in the infrared
(Hervig et al., 2001; Eremenko et al., 2005). However, in order to access
particle sizes efficiently, the deviation of the particle’s scattering/absorption
properties from the pure Rayleigh regime needs to be studied. This requires
wavelengths of the order of the particle size and, hence, measurements in
the visible and, in particular, the ultraviolet part of the spectrum. Previous
NLC measurements have been made at wavelengths down to 355 nm with
ground-based instruments (e.g. von Cossart et al., 1999) and down to about
200 nm with rockets and satellites (e.g. Rusch et al., 1991; Gumbel et al.,
2001). In paper IV we argue that measurements at even shorter wavelengths
in the extreme ultraviolet (EUV) are necessary to extend our knowledge of
the NLC particle size distribution. In particular, Lyman-α radiation at
121.57 nm is well suited for such studies as NLC are well illuminated by
solar Lyman-α during daytime conditions. At 121.57 nm even small
mesospheric ice particles come well into the Mie scattering regime and start
to influence the total NLC scattering. Therefore, an investigation of the
scattering and extinction properties of ice particles at Lyman-α can yield
important complementary information about the polar mesospheric ice
layer.
16
The SLAM instrument (Scattered Lyman-Alpha in the Mesosphere)
presented in paper IV is a new development and a test of the measurement
concept for NLC scattering of solar Lyman-α. The instrument is designed
as a double monochromator. A first flight opportunity for measuring
Lyman-α in the vicinity of NLC was in July 2006 with SLAM as part of the
eARI HotPay-1 sounding rocket from Norway. The SLAM instrument was
launched together with four other scientific instruments on 1 July 2006.
While this test flight was intended to be the first measurement of the
altitude- and angle-dependent scattering of Lyman-α by NLC in the
mesosphere, a malfunction of the payload shortly after launch unfortunately
made it impossible to collect any mesospheric data. A second opportunity
to measure the scattering of Lyman-α by NLC will be on the PHOCUS
sounding rocket (section 3).
1.2.2 Meteoric smoke
The Earth’s atmosphere is constantly bombarded by meteoric material.
Most of the incoming meteoroid dust is in the size range 10-300 µm
[Ceplecha et al., 1998]. Depending on entry speed and inclination, smaller
particles (< 20 µm) may survive almost intact as micrometeorites, while
larger particles (>100 µm) are believed to ablate and vaporize entirely in
the 70-130 km altitude region. Estimates of the total amount of incoming
material are still uncertain and vary between 10 and 100 tons per day [Love
and Brownlee, 1993; Mathews et al., 2001; von Zahn, 2005]. Also our
knowledge about the fate of the vaporized material is rather limited. It is
well recognized that meteoroids are the source of the metal atom layers that
are observed by lidar [Kane and Gardner, 1993] and satellite [Gumbel et
al., 2007], and much progress has been made in understanding the
chemistry of these metals [Plane, 2003]. It is further thought that this
vapour recondenses into secondary nanometre-size particles, which are then
subject to coagulation, sedimentation and global transport [Rosinski and
Snow, 1961; Hunten et al., 1980; Gabrielli at al., 2004; Megner et al.,
2006]. Processes of relevance for this “meteoric smoke” are schematically
summarized in figure 4.
17
Figure 4. Schematic of the fate of meteoric material in the atmosphere.
Once formed from the ablated meteoric material, meteoric smoke particles
are thought to play a major role in mesospheric processes. Examples are:
•
Smoke particles are today recognized as the major candidate for
condensation nuclei for ice particles and, hence, phenomena like
noctilucent clouds and polar mesosphere summer echoes (section
1.2.1).
•
Heterogeneous chemistry has been suggested to take place on the
surface of smoke particles. An example is the proposed catalytic
recombination of O and H2 as a local source of water in the
mesosphere [Summers and Siskind, 1999]. By scavenging various
gas-phase products of meteoric ablation, smoke is also thought to
act as (ultimate) sink in mesospheric metal chemistry [Plane,
2004].
•
Smoke can play a significant role in the D-region charge balance
[Rapp and Lübken, 2001]. While the particles efficiently scavenge
free electrons and positive ions, the charging may in turn strongly
affect smoke coagulation processes.
In the stratosphere meteoric particles are thought to constitute an important
part of the background aerosol. Murphy et al. [1998] showed that a high
proportion of stratospheric aerosol is meteoric. The importance of meteoric
smoke in the middle atmosphere relates to the fact that input from the
troposphere is small. The only dominant dust sources from below are major
volcanic eruptions and gas-to-particle conversion, mainly of sulphur
18
compounds complemented by the existence polar stratospheric clouds
(PSC), featuring varying composition (water ice, supercooled ternary
solutions (STS), nitric acid trihydrate (NAT)) and important effects on
ozone chemistry. Due to the strong downward circulation in the winter
polar vortex, meteoric smoke can be efficiently transported to the lower
winter stratosphere [Megner et al., 2008b]. Here they become involved in
the microphysics of PSC. In order to explain the observed composition of
PSC particles, Voigt et al. [2005] have recently suggested meteoric smoke
as a trigger of heterogeneous nucleation of NAT. Smoke particles have also
been used as a tracer for atmospheric circulation [Gabrielli et al., 2004;
Lanci and Kent, 2006].
Despite this potential importance of mesospheric particles very little is
known about their properties. This is because the only means of accessing
them is by in situ measurements from sounding rockets. These
measurements are very difficult and there are a number of challenges
inherent to sounding rocket experiments that need to be considered, such as
aerodynamics [Gumbel, 2001; Hedin et al., 2005] and charging processes
(payload charging). The basic aerodynamic challenge lies in the size of the
meteoric smoke particles. Smoke particles are so small that they tend to
follow the airflow around the payload rather than reaching the detector.
Since we want the particles to hit the detector surface, careful aerodynamic
design is of critical importance for smoke particle experiments. The
interpretation of the particle measurements requires also a detailed
understanding of the specific detector response, and this is far from trivial.
Basic instrumental questions are:
•
How are measured particle concentrations related to the
undisturbed particle concentrations in the atmosphere?
•
How are properties of detected particles related to the particle
properties in the atmosphere?
•
How is the charge measured by particle detectors related to the
charge of particles in the atmosphere?
Paper I and II focus on the aerodynamical aspects of these questions.
Measurement of charged smoke particles
Most data on meteoric smoke available today is indirect evidence for the
existence of the particles from measurements of heavy charge carriers.
19
Recent examples of charged particle measurements have been reported by
Gelinas et al. [1998], Horányi et al. [2000], Lynch et al. [2005], Rapp et al.
[2005, paper III], and Rapp and Strelnikova [2008]. The major class of
detectors aimed at the detection of charged meteoric smoke particles is
based on a detector design originally developed by Havnes et al. [1996] for
the study of ice particles in the polar summer mesosphere. The detector
concept uses a Faraday Cup for the detection of heavy charge carriers in
combination with biased grids that shield the electrode in the inner part of
the detector against contamination by electrons and light ions from the
ambient ionospheric D-region. Gelinas et al. [1998] were the first to apply
a detector of this kind to the rocket-borne study of meteoric smoke particles
and Lynch et al. [2005] further developed the design to allow the
discrimination between positive and negative particles.
Figure 5. Left panel: The top section of the ECOMA payload for mesospheric particle
studies. The ECOMA particle detector from IAP is in the centre and the MAGIC
particle sampler from MISU/NRL in the foreground. Other instruments are
temperature probes from IAP (left, right) and ionospheric instrumentation from the
Technical University Graz (booms). Right panel: The ECOMA detector.
By combining the original design of Havnes et al. [1996] with a xenon
flash lamp, Rapp et al. [2005, paper III] provided a way to detect neutral
particles by photoionization. This detector is the central part of the
ECOMA sounding rocket project (figure 5). The ECOMA project
(Existence and Charge state Of Meteor smoke particles middle
Atmosphere) comprises a series of 4 campaigns from Andøya Rocket
Range. The project is led by the Leibniz Institute of Atmospheric Physics
(IAP) and the Norwegian Defence Research Establishment (FFI). One
campaign in dark autumn conditions in September 2006 and two in polar
summer conditions in August 2007 and July 2008, respectively, have
already taken place. Results are described by Rapp and Strelnikova [2008]
and in paper III. Additional instruments onboard the payload concern
temperature, turbulence, ionospheric parameters, optical NLC properties
(MISU) and particle sampling (see below), thus providing a very
20
comprehensive characterisation of neutral and ionospheric conditions. The
campaigns are also supported by the extensive ground-based equipment
located at ALOMAR and Andøya Rocket Range, in particular the ALWIN
VHF radar, the SAURA MF wind radar, the SKiYMET meteor radar, and
the ALOMAR RMR and sodium lidar. In addition it is supported by the
EISCAT incoherent scatter radar facility.
In addition to Faraday cup type detectors other techniques have been
applied in the study of meteoric smoke. These include mass spectrometry
[Schulte and Arnold, 1992], Gerdien condenser [Croskey et al., 2001] and
magnetically shielded probes [Robertson et al., 2004; Smiley et al., 2006].
Signatures of charged particles have recently also been reported from
incoherent scatter radar measurements [Rapp, et al., 2007a]. It is however
important to note that none of the above measurements of heavy charge
carriers provides a definite proof that the detected species really are smoke
particles of meteoric origin. In order to provide such a proof and more
detailed studies of the particles, instruments are needed that directly sample
the smoke and this is the idea of the MAGIC meteoric smoke samplers
discussed next and in paper II.
Paper I analyses the aerodynamics of smoke particle measurements by
Faraday cup detectors. The paper provides a response function that
specifies the fraction of atmospheric particles that is actually detected by a
given instrument design as a function of particle size and charge, altitude,
and flow conditions.
Rocket-borne sampling of smoke particles
Because of the lack of knowledge about meteoric smoke particles, they
have less seriously been called “magic dust”. This is the origin of the name
of the MAGIC rocket project (MAGIC: Mesospheric Aerosol - Genesis,
Interaction and Composition). The major objective of the MAGIC project
was the development of a sampling instrument and analysis techniques that
for the first time would allow smoke particles to be brought back from the
mesosphere to the laboratory and to study their properties in detail. The
MAGIC detector concept has been developed in collaboration between the
Naval Research Laboratory in Washington D.C. and German and Swedish
scientists [Gumbel et al., 2005]. As nanometre-size particles tend to follow
the airflow around payload structures, a basic idea of MAGIC is to
minimise aerodynamical perturbations in the measurement volume so that
the particles could be sampled. This resulted in an instrument design with
sampling surfaces of 3 mm diameter located on top of pins that are ejected
21
from the payload top (see figure 6). The aerodynamics of the MAGIC
meteoric smoke sampler is the topic of paper II. The sampling surfaces are
standard transmission electron microscopy (TEM) grids and can be taken
directly from the measurement pin to the microscope for analysis. The basic
scientific questions addressed in this project are:
•
Do re-condensed smoke particles of cosmic origin exist in the
mesosphere?
•
What is their number density and size distribution?
•
What is the spatial distribution of the particles and how are they
transported?
•
What is the elemental and molecular composition of the particles?
•
What is their charge state and how do they interact with their
mesospheric and ionospheric environment?
Figure 6. a) The MAGIC collection unit with extended sampling pin. b) The top of
MAGIC sounding rocket payload with a sampling probe extended from one of the
three MAGIC sampling units surrounding the Hygrosonde instrument [Khaplanov et
al., 1996; Lossow et al, 2008a] in the centre. c) TEM grid mounted with copper cap on
top of an extended sampling pin.
Earlier rocket campaigns with the MAGIC instruments were conducted
from Esrange, Sweden, in January 2005 [Gumbel et al., 2005], from
Wallops Island, USA, in May 2005, and as part of the German/Norwegian
ECOMA project from Andøya, Norway, in September 2006, August 2007
and July 2008. However, these flights have not yet resulted in any
quantitative results. The September 2006 campaign featured launches of
two payloads, each carrying two MAGIC’s. Analysis of the first launch did
not give quantitative results due to contamination of the sampling grids,
most likely from the payload itself, while the second payload was lost due
22
to problems with the recovery system. Recovery problems also led to the
lack of results in 2007. In July 2008 three payloads were successfully
launched, this time with one MAGIC particle sampler on each payload.
Samples from these flights are currently being analysed, but quantitative
results are not yet available.
The analysis of the MAGIC smoke samples after recovery is primarily
based on detailed TEM studies at Naval Research Laboratory and, since
2008, at a new Centre for advanced electron microscopy studies in the
Department of Physical, Organic, Inorganic and Structural Chemistry at
Stockholm University. The TEM analysis is intended to provide particle
numbers, sizes and shapes, while the combination of the TEM technique
with energy-filtered imaging and energy-dispersive x-ray spectroscopy can
provide the elemental composition of the sampled particles.
MAGIC is currently scheduled for one future rocket flight as a central part
of the PHOCUS project, focusing on the interaction between meteoric
smoke, noctilucent clouds and gas-phase chemistry (section 3).
1.2.3 Nightglow
The atmosphere is constantly emitting radiation. This radiation called
airglow is emitted by excited atoms, molecules or ions generated in a
variety of production and excitation processes following the interaction of
sunlight with atmospheric constituents during the day. The airglow can be
divided into several types depending on the time of day they are observed
and include dayglow, twilightglow and nightglow. The nightglow is rather
faint and can be compared in intensity to the light of a candle at a distance
of 100 m. Although the dayglow intensities are several orders of magnitude
larger, it is not detectable by eye because of the strong scattering of sunlight
in the atmosphere. The airglow is driven by photons from the sun, whereas
auroras are excited by the impact of energetic particles. Outside the auroral
region or in times of quiet geomagnetic conditions, the terrestrial night sky
spectrum is dominated by emission features related to oxygen
photochemistry. Molecular oxygen has relatively low dissociation energy
and O2 photolysis, mainly in the Schuman-Runge bands and continuum, is
therefore a very efficient way of converting solar radiation into chemical
energy.
The energy is deposited during daytime and the part not directly re-radiated
or converted into heat is stored in the form of excitation of metastables and
23
dissociation energy. Atomic oxygen is the major carrier of this chemical
energy in the upper atmosphere and it is the recombination of atomic
oxygen that generates molecular oxygen in a number of metastable states.
With the energy available in this recombination reaction, high vibrational
levels of the ground state (X 3Σ −g ) or any of the five bound states (a1∆ g ,
b1Σ +g , c1Σ u− , A3Σ u+ , A'3 ∆ u may be formed. Another weakly bound state, 5∏g,
)
has been theoretically predicted but not yet observed and has been proposed
as a likely precursor to the above lower-lying light-emitting states. Through
a number of energy transfer and quenching processes the excited O2 then
gives rise to nightglow emissions covering the spectral range from the
ultraviolet to the infrared [Meriwether, 1989; Mlynczak, 1999; Wayne,
2000]. These processes cause a complex and strongly altitude-dependent
photochemistry between about 85 and 105 km. Other emissions that
contribute to the nightglow are emissions from atomic oxygen [Barth and
Hildebrandt, 1961], and vibrationally excited OH molecules [Bates and
Nicolet, 1950]. A fourth source of nightglow emissions is the emissions of
metallic atoms such as sodium, calcium, potassium and magnesium. Also
for these emissions, atomic oxygen is essential as its reaction with metal
oxides is the major production channel of metal atoms in excited states
[Plane, 2003].
Although it is generally accepted that the energy of excitation originates
from the recombination of oxygen atoms, the exact mechanisms and
reaction rates took long time to be established, and some have still not been
established today [Slanger and Copeland, 2003]. Much of today’s
knowledge is based on the Energy Transfer in the Oxygen Nightglow
(ETON) campaign that was conducted in March 1982 from South Uist, U.
K. [Greer et al., 1986]. A series of seven payloads was launched to
measure altitude profiles of emission features and atmospheric
composition. There have been several rocket experiments in the past with
the objective to simultaneously study atomic oxygen density and oxygen
airglow intensity [e.g. Thomas, 1981; Sharp, 1980; McDade et al., 1984].
However, the most extensive of these multi-instrumented rocket studies
was the set of campaigns with OXYGEN in March 1981 [Witt et al., 1984]
followed by the ETON venture. The analysis of the ETON database
resulted in detailed knowledge about nightglow excitation and quenching
mechanisms [Greer et al., 1986; McDade et al., 1986a; 1986b; 1987a;
1987b; Murtagh et al., 1986]. This was not only interesting for atmospheric
scientists but brought new understanding also to the basic physics of
excited molecular states. In this sense, ETON used the mesosphere and
lower thermosphere as a laboratory that provided experimental conditions
that could hardly be achieved in conventional laboratory experiments.
24
Today, basic studies of airglow mechanisms are still an active field of
research. But more importantly, airglow emissions are today applied as a
versatile tool for atmospheric research. Remote sensing of airglow
emissions from the ground or space is being used for studies of middle
atmospheric composition, structure and dynamics over a wide range of
spatial and temporal scales [e.g. Mlynczak et al., 2000; French et al., 2005;
Taylor et al., 1997; Degenstein et al., 2005]. The rocket studies originally
relating the airglow processes to the state of the atmosphere have been the
basis for these techniques. Measurements of the emission rates yield
information about atmospheric composition, while studies of the Doppler
shifts and the ratios of the airglow lines provide information about winds
and other motions, and temperature, respectively [e.g. Shepherd et al.,
1993; Espy et al., 1997]. Measurements of line shifts and shapes are mainly
performed from satellites or from ground-based instruments while the
emission rates are readily measured also from sounding rockets. Groundbased imaging of nightglow patterns is used to study dynamics in terms of
gravity wave activity in the upper mesosphere [e.g. Taylor et al., 1997].
Airglow emissions have also been shown to be fundamental measures of
energy deposition from which rates of atmospheric heating can readily be
derived [Mlynczak, 1999].
In paper V we present another use of O2 airglow emissions. We
demonstrate how airglow measurements can be used to calibrate direct
atomic oxygen measurements by the resonance fluorescence technique
from sounding rockets. The resonance fluorescence technique is sensitive
and provides good altitude resolution, but is in itself notoriously difficult to
calibrate [Gumbel et al., 1998; Dickinson et al., 1980; Sharp, 1991]. As
atomic oxygen is the major carrier of chemical energy in the MLT region
accurate knowledge about the vertical distribution of O is crucial for any
comprehensive study of chemical and dynamical processes in this part of
the atmosphere [Ward and Fomichev, 1993; Mlynczak and Solomon, 1993].
25
Chapter 2
Atmospheric science with sounding
rockets
2.1 The need for sounding rockets
Sounding rockets are crucial as tools for atmospheric research. It is the
features of the mesosphere and lower thermosphere that define the need for
sounding rockets. As mentioned earlier, complex interactions between
radiative, dynamical, microphysical and chemical processes give rise to
many prominent phenomena such as NLC, PMSE, the ablation and
transformation of meteoric material, or the Earth’s nightglow. Interactions
on small spatial scales are a common feature of these phenomena and it is
here that fundamental physics and new science can be revealed in the
mesosphere today. When aiming at an understanding of these processes, we
need to address interacting parameters simultaneously and on the small
spatial scales of interest. Rocket-borne in situ studies with dedicated sets of
experiments meet both these requirements. This cannot be provided by
remote sensing techniques from the ground or from space.
For half a century sounding rocket experiments have thus been crucial as a
source of knowledge about the mesosphere and thermosphere. In Sweden,
the first scientific rocket was launched in 1961. Five years later Esrange
was established. In Norway, the first scientific rocket was launched in 1962
from Andøya Rocket Range. The complexity of scientific questions has
generally increased throughout the history of these rocket projects. In the
early years, a new altitude profile of a single atmospheric parameter was
certainly well worth a publication. Today, the strength of rocket projects
26
often lies in the study of complex interactions that involve many
atmospheric parameters. In many cases, this increase in complexity has
been based on a development towards larger multi-instrument payloads.
But also the combination of in situ and remote sensing techniques has
contributed to the growing complexity of rocket-borne studies.
Sounding rockets keep their important role in a world of global satellite
measurements. Scientifically, the relationship between rockets and satellites
is defined by the different nature of their measurements. The local
'snapshot' measurement makes sounding rockets a natural complement to
the large-scale data provided by satellites. This can be used in different
ways such as validation of satellite measurements by sounding rockets or
real co-analysis of local and large-scale datasets. In addition to direct data
validation and co-analysis, satellite research and rocket research have been
inspiring one another in many other and fascinating ways. For example, in
situ and ground-based results from the MaCWAVE/MIDAS rocket
campaign indicated that unusual planetary wave activity in the southern
hemisphere winter stratosphere could control the thermal conditions in the
northern hemisphere summer mesosphere [Goldberg et al., 2004].
Subsequently, the existence of such an interhemispheric coupling was also
suggested by general circulation modelling [Becker and Fritts, 2006].
Based on these ideas, global noctilucent cloud data from the Odin satellite
were then analysed and, indeed, showed a significant interhemispheric
coupling on a year-to-year basis [Karlsson et al, 2007]. This is one of many
examples of rockets and satellites jointly contributing to the development of
new scientific ideas.
Last but not least, an important connection between rocket and satellite
projects is the exchange of technology. Many of today's remote sensing
instruments on satellites build on experience from many years of rocketborne optical instruments. Also the opposite way is possible: The
experience with the submillimetre/millimetre radiometer onboard the Odin
satellite [Murtagh et al., 2002] is today the basis for a new type of rocketborne water vapour radiometer proposed for the PHOCUS rocket project
(section 3).
2.2 Challenges
In situ experiments in the mesosphere are challenging. Many scientific and
technical requirements need to be fulfilled by sounding rocket instruments.
27
The scientific skills behind rocket instruments and the need for engineering
experience and continuity are not to be underestimated.
Scientifically, a basic requirement is to obtain quantitative measurements.
This is far from trivial. The obstacles on the way from a simple detection
idea to a quantitative rocket experiment may in the best case be surprisingly
many, in the worst case they are just impossible to overcome. The first
steps involve the physical measurement principle, the instrument design,
and the design of appropriate calibration techniques. Beyond these steps
come the challenges of the actual rocket flight. Most notably, our
understanding of interactions between the rocket payload and its neutral
and charged environment is still too limited. After half a century of
sounding rockets, there remain a remarkable number of unsolved questions.
New experimental as well as theoretical approaches are required to improve
this situation.
An instructive example is the history of rocket-borne measurements of
atomic oxygen [Gumbel, 1999]. Resonance fluorescence has been used by
many groups and calibration techniques have been developed in the
laboratory, in situ and by theoretical means [e.g. Dickinson et al., 1980;
Sharp, 1991; Gumbel and Witt, 1997]. Nevertheless, measurement results
show a spread that is most likely not geophysical. Paper V discusses these
issues further.
Also the particle measurements discussed in chapter 1.2.2 can illustrate
these points. What particle sizes can we actually detect? A basic
aerodynamic challenge lies in the fact that nanometre-size particles tend to
follow the gas flow around the payload structure rather than reaching the
detector. This leads to a size-dependent and altitude-dependent detection
efficiency that needs to be quantified [papers I and II]. How do we actually
measure charged particles? Many in situ measurements of mesospheric
particles have been limited to the charged fraction of the particle population
[Rapp et al., 2007a; paper III]. However, it has recently been speculated
that impact charge measurements may be strongly influenced by secondary
charging effects like fragmentation or triboelectric charging [Barjataya et
al., 2006, Havnes and Naesheim 2007]. What processes influence the
electric potential of the rocket payload? Payload charging can strongly
influence measurements of ionospheric properties or particles [Holzworth et
al., 2001]. Payload charging mechanisms are today well understood when it
comes to satellite platforms at higher altitudes. However, the intimate
interplay of aerodynamics and charge flows makes the problem much more
complicated for sounding rockets. Charging effects have puzzled scientists
after many rocket flights and we are today still far from the necessary
28
answers. An important step has recently been taken by combining
numerical simulations of aerodynamics, charge flows and electric fields
near a sounding rocket [Sternovsky et al., 2004]. But the complete picture is
complex and more comprehensive efforts will be necessary to
quantitatively understand the interaction of rocket payloads with their
environment.
29
Chapter 3
The PHOCUS project
Aerosol particles are currently the hottest topic in mesospheric research.
Particle species comprise ice particles, smoke particles of meteoric origin,
and possibly other background particles formed by conversion from trace
gases. Important questions concern both the properties of particle layers
and their interaction with various phenomena in the mesosphere and lower
thermosphere. This includes the relationship between smoke and ice, ice
particle nucleation and evolution, and the possible influence of these
particle species on neutral and ion chemistry. The scientific importance of
mesospheric aerosols is reflected by the fact that there are currently two
major rocket efforts in Europe focussing on this topic. One is the
German/Norwegian ECOMA programme from Andøya Rocket Range
(section 1.2.2), the other is the Swedish PHOCUS project (Particles,
Hydrogen and Oxygen Chemistry in the Upper Summer mesosphere) from
Esrange which is the topic for this chapter. Both aim at the characterisation
of smoke and ice particles and their interactions. While ECOMA
concentrates on interactions with ionospheric processes, PHOCUS
concentrates on interactions with neutral chemistry. Important for the
ECOMA and PHOCUS projects is the fact that most of the scientific groups
involved are contributing with instruments to both projects.
PHOCUS is funded by the Swedish National Space Board and the sounding
rocket campaign is scheduled for a launch from Esrange, Sweden, in the
summer of 2010. The results of this thesis provide important input to the
preparation and the scientific analysis of the project. In the following the
science questions and details about the individual instruments on board
PHOCUS and their development will be discussed.
30
3.1 Science questions
With PHOCUS we intend to study the interaction between three important
mesospheric players that earlier have been addressed independently of each
other: meteoric smoke, noctilucent clouds, and odd oxygen/odd hydrogen
chemistry. The PHOCUS project brings three lines of scientific
developments at MISU together. This includes the optical probing of
mesospheric chemistry, the optical characterization of NLC particles, and
the study of mesospheric smoke particles within the MAGIC programme.
This is complemented by our partners focusing on particle/ionosphere
interactions and the distribution of water vapour. The basic science
questions related to these particle interactions are summarized in Figure 7.
Figure 7. Schematic of the science questions addressed by PHOCUS.
Smoke and NLC
The direct connection between NLC and their potential condensation nuclei
has never been investigated experimentally. This is a critical shortcoming
for our current understanding of NLC processes. Meteoric smoke particles
have been suggested as dominant condensation nuclei for ice particles [e.g.,
Rapp and Thomas, 2006; Plane, 2000]. The role of ions and charged
particles in ice nucleation remains unclear [Gumbel and Witt, 2002;
Gumbel and Megner, 2008]. Recently, also sulphate compounds have been
suggested as a new actor in mesospheric particle processes [Mills et al.,
31
2005]. Today we neither know typical number densities nor particles sizes
of various aerosol types in the summer mesosphere. Both NLC nucleation
rates and ice particle properties depend strongly on these parameters [e.g.,
Turco et al., 1982]. Important open questions also concern the shape of
NLC particles, with growing evidence that particles are not spherical [Rapp
et al., 2007b; Eremenko et al., 2005].
NLC and HOx
Earlier rocket-borne measurements of atomic oxygen and NLC showed
distinct O minima near the cloud altitude, leading to the idea that ice
particles might influence mesospheric chemistry [Kopp et al., 1985;
Gumbel et al., 1998]. These questions have recently been addressed by
laboratory experiments [Murray and Plane, 2005], showing that
heterogeneous processes on NLC surfaces are not efficient enough to
explain the observed oxygen loss. Rather, NLC particles can influence
atomic oxygen indirectly by affecting the distribution of water vapour and,
thus, gas-phase HOx chemistry. A shortcoming of earlier in situ NLC
studies is that all rocket flights were carried out during twilight when the
solar Lyman-α flux exhibits a strong gradient exactly at NLC altitudes.
This has a strong influence on the altitude-dependent chemistry and makes
it difficult to distinguish particle effects from radiation effects. This is one
strong reason why the PHOCUS payload should be launched during midday rather than in twilight.
Smoke and HOx
Influences of particle layers on mesospheric chemistry have also been
discussed in connection with meteoric smoke particles [Summers and
Siskind, 1999]. Recombination of H2 and O on these particles has been
suggested to explain observations of mesospheric HOx chemistry and
unexpectedly high water concentrations. A conclusion from these
observations has been that the odd hydrogen/odd oxygen chemistry in the
middle atmosphere is not well understood and might not be explainable
with conventional gas-phase chemistry alone. This problem has been
denoted as the "HOx dilemma" [Conway et al., 2000].
Water vapour
A fourth atmospheric player intimately connected to the above three is
water vapour. A sufficient supply of water vapour is the basic pre-condition
for the formation of ice particles. The availability of water vapour at the
summer mesopause is determined by a delicate balance with upward
32
transport as the source and photolysis in 24-hour sunlit conditions as an
efficient sink. At the same time, being the source of odd hydrogen species
(HOx), water vapour together with atomic oxygen controls the chemical
environment in the mesopause region. Large effects of NLC on the (re-)
distribution of water vapour have earlier been reported by Summers et al.
[2001] and modelled e.g. by von Zahn and Berger [2002]. With Odin's
measurements, we are now able to study the distribution of water vapour on
a global scale [Lossow et al., 2008b]. However, in situ water vapour
measurements in the daytime mesosphere have so far not been available.
With the recent Swedish advances in space-borne microwave technology
for the Odin satellite and related projects, such measurements are now
feasible as a part of PHOCUS.
Temperature
The existence of ice particles in the mesopause region is a direct
consequence of the extremely low temperatures in the altitude range
between 80 and 90 km. On top of the mean temperature structure, local
deviations occur due to gravity. Rapp et al. [2002] showed a substantial
influence of this gravity wave activity on the formation of NLC. Hence, in
situ temperature measurements with high spatial resolution are needed
together with accurate measurements of water vapour when NLC are to be
analysed in terms of ice condensation and sublimation processes.
Ionospheric parameters
Charging processes are closely connected to a variety of layered particle
phenomena in the mesosphere. The most prominent example of this
interaction is the existence of polar mesosphere summer echoes [Rapp and
Lübken, 2004]. In fact, for typical ionospheric conditions, efficient negative
particle charging can be expected due to capture of electrons [Rapp and
Lübken, 2001]. Even positive particle populations have been observed
[Havnes et al., 1996; paper III], although so far in lack of a conclusive
explanation. Interactions between ion chemistry and the existence of
particles are mutual: Particles influence ion composition and reaction
schemes by capture of electrons and ions. Charges captured by particles, on
the other hand, can influence further particle growth by coagulation
processes [Jensen and Thomas, 1991]. In summary, a comprehensive
characterisation of particle processes in the mesosphere is not possible
without a simultaneous analysis of the state of the background ionosphere.
In addition to the in situ measurements planned for the PHOCUS rocket
payload, the objectives described above require complementary ground33
based data (NLC and PMSE monitoring) as well as detailed laboratory and
model studies. The schematic overview in Figure 8 summarizes the various
parts of the mission. Details about the individual instruments and their
development are discussed in the next section.
Figure 8. A schematic overview over the in situ, ground-based, laboratory and
model efforts contributing to the PHOCUS project from Esrange in 2010.
Table 1. Instrumentation for the PHOCUS sounding rocket.
Instrument
Experimenter
Objective
O-probe
MISU
atomic oxygen density
H-probe
MISU
atomic hydrogen density
2× IR photometers
MISU
oxygen and hydrogen related airglow
MAGIC
MISU/NRL
smoke particle properties
2× polarisation photometers
MISU
NLC particle properties
phase function photometer
MISU
NLC particle properties
2× microwave radiometer
Chalmers
water vapour density
2× Faraday cups
IAP and UiT
charged dust particles
CDD
LASP
charged dust particles
CONE
IAP/FFI
temperature, density, turbulence
Faraday, ion probe
TU Graz
electron and ion densities
34
3.2 Instrumentation
The backbone of the PHOCUS rocket campaign is an instrumented payload
for in situ measurements of particles, atomic and molecular species as well
as atmospheric background conditions. Table 1 summarizes the proposed
instrumentation. All experiment concepts are discussed briefly below.
Atomic oxygen probe: The atomic oxygen measurement is based on
MISU’s experience with rocket-borne composition studies by VUV
resonance techniques [Gumbel and Witt, 1997]. Resonance fluorescence of
the 3S-3P triplet at 130.6 nm is used for the detection of the O profile in the
mesosphere and lower thermosphere. We now plan the development of a
smaller probe, focusing entirely on the sensitive resonance fluorescence
technique. This is a pre-condition for obtaining a light-weight and costeffective instrument for small rocket payloads. The direct measurement will
be calibrated by a simultaneous photometric dayglow measurement of the
(0-0) band of the O2 (a1∆ g → X 3Σ −g ) IR Atmospheric system at 1.27 µm as
described in paper V. This dayglow emission is related to the photolysis of
ozone, which during the day is in steady state with atomic oxygen and a
retrieval of atomic oxygen is possible [Mlynczak et al., 1993]. Other
laboratory or in situ techniques for calibrating resonance fluorescence
payloads are in principle available, but they are complicated and expensive
and they exhibit inherent uncertainties [Dickinson et al., 1980; Sharp, 1991;
Gumbel et al., 1998]. With the photometric calibration proposed in paper
V, many of these problems are avoided.
Atomic hydrogen and Lyman-α probe: The atomic hydrogen
measurement uses a similar resonance fluorescence technique, employing
the 121.6 nm H Lyman-α line [Sharp and Kita, 1987]. This technique has
been used by MISU during the NLC-91 campaign. Data from these
measurements have been analysed. However, due to sensitivity of the
detector to background radiation at longer wavelengths no useful results
could be obtained at that time. The detector developments and
measurements of the diffuse daytime Lyman-α fluxes described in paper IV
were intended as an important input to the preparation of the PHOCUS
project. Similar to the atomic oxygen measurements, also the direct
measurement of atomic hydrogen will be calibrated by photometric airglow
measurements. This will be based on photometry of OH Meinel emissions
in the near infrared. As the OH Meinel emission is produced by the reaction
35
between mesospheric O3 and H, and O3 concentrations are available from
the O2 IR Atmospheric band photometry (paper V), H concentrations can
be deduced. In addition to the resonance fluorescence of H atoms
[Meriwether, 1989], the detector on PHOCUS will also measure solar
Lyman-α scattered by NLC particles. As discussed in section 1.2.1, this
will complement the longer wavelength scattering measurements by the
PHOCUS polarisation photometers and the phase function photometer
described below.
NLC photometry: NLC photometry is currently the best technique
available for determining altitude ranges of NLC in situ. At the same time,
UV photometry allows a study of particle properties like size and shape by
analysing the spectral dependence, phase function, and polarisation of the
scattering [Witt, 1969; Gumbel et al., 2001]. On PHOCUS, we will fly a
unique photometer package that for the first time will investigate all three
parameters simultaneously. Two forward-viewing photometers will
measure at different wavelengths and will both be equipped with polarisors.
A third photometer will be tilted for viewing the sky under varying
scattering angles. The design is based on photometer developments earlier
flown on the DROPPS and ECOMA/MASS rocket campaigns [Gumbel et
al., 2001; Megner et al., 2008c].
The polarisation photometers will be equipped with fixed linear polarisors.
They will be mounted in forward direction with their field of view parallel
to the rocket spin axis. In order to gain maximum information,
measurements are needed in the Mie regime and, hence, one photometer
will measure in the UV at 220 nm. The second will be centred at 450 nm in
order to study the spectral dependence. The payload spin will be utilised to
scan through the polarisation direction, thus providing us with the Stokes
vectors I, Q and U at both wavelengths. The phase function photometer will
be mounted sideways, viewing the overhead sky at an angle of 30° from the
rocket spin axis. Due to the payload spin, NLC will be observed under
varying scattering geometries as the payload approaches the cloud layer
[Gumbel et al., 2001]. In addition, as mentioned above, the detector for the
H resonance fluorescence measurement will also measure the Lyman-α
light at 121.6 nm scattered by ice particles thus providing NLC
measurements at different scattering angles and at even shorter
wavelengths.
Particle impact measurements: As a large fraction of mesospheric
particles can be expected to be charged, charge-sensitive detectors can
provide information about particles and their distribution (section 1.2.2).
The two Faraday cup detectors contributed by IAP and the University of
36
Tromsø (UiT) are based on the original design by Havnes et al. [1996] and
will measure heavy charged particles. Discrimination against light ions is
provided by electrostatic shielding. In addition, the UiT detector is designed
as to distinguish the generation of secondary charges upon ice particle on
the detector surface [Havnes and Naesheim, 2007]. Faraday cup
measurements are discussed further in paper I and III.
Also the Colorado dust detectors (CDD) provided by the Laboratory of
Atmospheric and Space Physics (LASP) will measure charged particle
concentrations. They are graphite patch collectors mounted flush with the
skin of the payload and measure a net current deposited on the detectors by
heavy aerosol particles [e.g. Horányi et al, 2000; Amyx et al., 2008]. For
the CDD’s discrimination against light ions is provided by magnetic
shielding.
Particle sampling: Direct sampling of meteoric smoke particles will be
made with the MAGIC meteoric smoke sampler (section 1.2.2). The aim is
a comprehensive characterisation of nanometre-size meteoric particles.
With a new facility for electron microscopy at Stockholm University, good
opportunities for the laboratory analysis of the PHOCUS/MAGIC results
are now available in Sweden. The sampling of meteoric smoke particles by
the MAGIC meteoric smoke sampler is discussed further in paper II.
Water vapour measurements: The concept of a rocket-borne microwave
radiometer for water vapour is being developed by Chalmers University of
Technology. Based on substantial recent advances in microwave
technology [Murtagh et al., 2002], the measurement will be based on the
183 GHz and 558 GHz water lines. It is today fully feasible to build a
suitable mm-wave instrument well within the size of a rocket module and
without the need for active cooling. In fact, long-going design work for
such an instrument is already available at Chalmers in preparation of the
Atacama Large Millimeter Array (ALMA). The instrument for PHOCUS
will utilise one forward-looking antenna, thus detecting water vapour
emissions from the atmospheric column above the payload, and one sidelooking antenna for high altitude resolution. It is interesting to note that
already in the early 1990s a similar concept for mesospheric water
measurements was proposed by Croskey et al. [1993]. However, only now
has microwave technology advanced sufficiently to make an easy
incorporation of such an instrument into a sounding rocket feasible. The
perspective of obtaining local measurements of water vapour from the polar
summer mesosphere is extremely attractive to the entire middle atmosphere
science community.
37
Temperature measurements: The Combined Neutral and Electron sensor
(CONE) is contributed by the Leibniz-Institute for Atmospheric Physics
(IAP) and the Norwegian Defence Research Establishment (FFI). Basically,
CONE is an ionisation gauge for the neutral density measurements.
Temperatures are accurately derived from the neutral density via the
hydrostatic equation. CONE is today the only instrument available that can
resolve temperature structures in the mesopause region down to the level of
gravity waves [Rapp et al., 2002]. It also provides measurements of
turbulence on scales down to 10 cm [Lübken et al., 2002]. CONE has been
flown before from Esrange on the NLTE payloads in 1998 [Lübken et al.,
1999].
Ionospheric conditions: Ionospheric instrumentation will be provided by
the Technical University Graz. The instrument package consists of a
Faraday Radio Wave Propagation experiment (RWP) for high-accuracy
electron density measurements and an ion probe for ion densities. The RWP
uses four radio waves in the HF-range (1.3 to 7.8 MHz) that are transmitted
from the ground to the flying rocket payload. Electron densities are inferred
from the change of the polarisation without influences of payload charging
or aerodynamics [Jacobsen and Friedrich, 1979]. Positive ions are
measured with good altitude resolution by a gridded sphere with a
positively biased collector inside [Thrane, 1986]. In combination, this
instrument package provides both absolute calibration and high
sensitivity/altitude resolution. By separately measuring number densities of
electron and light positive ions, the charge balance is studied and detailed
information on both ionospheric conditions and ion/particle interactions can
be obtained.
38
Chapter 4
Results of this thesis
Paper I: This paper concerns the aerodynamics of rocket-borne meteoric
smoke particle impact measurements. Basically, the small smoke particles
tend to follow the gas flow around the payload rather than reaching the
detector if aerodynamics is not considered carefully in the detector design.
Quantitative ways are needed that relate the measured particle population to
the atmospheric particle population. This requires in particular knowledge
about the size-dependent, altitude-dependent and charge-dependent
detection efficiency for a given instrument. In this paper, we investigate the
aerodynamics for a typical electrostatic detector design by first quantifying
the flow field of the background gas and then introduce particles in the flow
field to determine their trajectories around the payload structure. We use
two different models to trace particles in the flow field, a Continuous
motion model and a new Brownian motion model. Detection efficiencies
are determined for three detector designs, including two with ventilation
holes to allow airflow through the detector. Results from this investigation
show that Brownian motion is of basic importance for the smallest particles
as they can be completely decelerated by the stagnating airflow inside the
detector and their further motion is governed by Brownian diffusion.
Results also show that the detection efficiency for meteoric smoke particles
is very much size and altitude dependent for this kind of particle impact
measurements. Particles have to be larger than 1 - 2 nm in radius to be
detected at any altitude between 70 and 95 km. The altitude range 70 - 75
km is an aerodynamical lower detection limit for particles up to at least 5
nm. It is important to note that earlier measurements of sharp cut-offs in
particle concentrations at these altitudes have been interpreted as a real
geophysical feature rather than instrumental in nature. Providing venting
holes allowing air to flow through the detector can improve the detection
39
efficiency at lower altitudes. However, the flow has to be relatively large in
order to significantly improve the detection efficiency.
Paper II: This paper discusses rocket-borne sampling of neutral meteoric
smoke particles and the detection efficiency of the MAGIC meteoric smoke
samplers. The MAGIC project (Mesospheric Aerosol – Genesis, Interaction
and Composition) aims to quantitatively answer fundamental questions
about the properties of smoke particles in the atmosphere. This is pursued
with the development of an instrument and analysis techniques that for the
first time allow smoke particles to be brought back from the mesosphere
into the laboratory for a detailed study of their properties. The idea with the
MAGIC smoke samplers is to minimise the aerodynamic perturbations in
the measurement volume by reducing the detector size to the order of the
molecular mean free path. The sampling surfaces are 3 mm diameter carbon
grids for transmission electron microscopy (TEM) mounted on top of pins
that are extended from the payload top and reach outside the payload shock
front. When the samples are returned to ground these grids can be moved
directly into the TEM for analysis. In order to characterise the sampling
process, we have performed simulations of the trajectories of nanometresized smoke particles towards the MAGIC detectors with the Brownian
motion model described in paper I. As a result, we obtain the detection
efficiency for the MAGIC detectors as a function of altitude and particle
size. Our simulations confirm that particles of radii down to 0.75 nm impact
on the sampling surface with an efficiency exceeding 80% over the entire
mesospheric altitude range of interest.
Paper III: In this paper we present results of in situ measurements of
charged nanoparticles, electrons, and positive ions obtained during a
sounding rocket flight in October 2004 from Kiruna, Sweden, under
nighttime conditions. The particle measurement reveals positive charge
signatures in the altitude range between 80 and 90 km corresponding to
peak charge number densities of ~100 e/cm3 at around 86 km.
Aerodynamical analysis of the sampling efficiency of our instrument using
the Continuous motion model described in paper I reveals that the particles
must have been larger than 2 nm assuming spherical particles with a density
of 3 g/cm3. The plasma environment of the observed particles is dominated
by negative and positive ions, with only few free electrons. A calculation of
the mean particle charge expected for particles in a plasma consisting of
electrons and positive and negative ions shows that the presence of
sufficiently heavy and numerous negative ions (i.e., mn > 300 amu and λ ≥
50) can explain the observed positive particle charge.
40
Paper IV: This paper concerns the study of mesospheric ice particles by
optical methods. In order to better understand noctilucent clouds (NLC) and
their sensitivity to the variable environment of the polar mesosphere, more
needs to be learned about the actual cloud particle population. Optical
measurements are today the only means of obtaining information about the
size of mesospheric ice particles. In order to efficiently access particle
sizes, scattering experiments need to be performed in the Mie scattering
regime, thus requiring wavelengths of the order of the particle size. We
show that NLC measurements in the extreme ultraviolet, in particular using
solar Lyman-α radiation at 121.57 nm, are an efficient way to further
promote our understanding of NLC particle size distributions. Here, we
present examples from recent rocket-borne studies that demonstrate how
ambiguities in the size retrieval at longer wavelengths can be removed by
invoking Lyman-α. We discuss basic requirements and instrument concepts
for future rocket-borne NLC missions. In order for Lyman-α radiation to
reach NLC altitudes, high solar elevation and, hence, daytime conditions
are needed. Considering the effects of Lyman-α on NLC in general, we
argue that the traditional focus of rocket-borne NLC missions on twilight
conditions has limited our ability to study the full complexity of the
summer mesopause environment.
Under the 6th Framework Program of the European Union, the European
Commission has promoted the utilization of the Arctic Lidar Observatory
for Middle Atmosphere Research (ALOMAR) at Andøya Rocket Range,
Norway. In particular, the enhanced ALOMAR Research Infrastructure
(eARI) included two sounding rocket launches in 2006 and 2008. A rocketborne instrument to measure the scattering of Lyman-α by ice particles was
developed and launched on the first of these sounding rockets. The
instrument is a double monochromator using two 3600 lines/mm
holographic gratings to give a good selectivity of the Lyman-α line and
provide a blocking of other wavelengths by 6 - 8 orders of magnitude.
Unfortunately, a malfunction of the payload shortly after launch made it
impossible to test the instrument concept and collect any mesospheric data.
Paper V: In this paper the possibility to calibrate direct atomic oxygen
measurements by simultaneous airglow measurements is presented.
Accurate knowledge about the distribution of atomic oxygen is crucial for
many studies of the mesosphere and lower thermosphere. Direct
measurements of atomic oxygen by the resonance fluorescence technique at
130.6 nm have been made from many sounding rocket payloads in the past
41
with good sensitivity and altitude resolution. However, accuracy is a
problem as calibration and aerodynamics make the quantitative analysis
challenging. In general, accuracies better than a factor 2 are not to be
expected from direct atomic oxygen measurements. As an example, we
present results from the NLTE sounding rocket campaign at Esrange,
Sweden, in 1998, with simultaneous O2 airglow and O resonance
fluorescence measurements. O number densities are found to be consistent
with the nightglow analysis, but only within the uncertainty limits of the
resonance fluorescence technique. Based on these results, we describe how
better atomic oxygen number densities can be obtained by calibrating direct
techniques with complementary airglow photometer measurements and
detailed aerodynamic analysis. Night-time direct O measurements can be
complemented by photometric detection of the O2 (b1Σ) Atmospheric Band
at 762 nm, while during daytime the O2 (a1∆) Infrared Atmospheric Band at
1.27 µm can be used. The combination of a photometer and a rather simple
resonance fluorescence probe can provide atomic oxygen profiles with both
good accuracy and good height resolution.
42
Chapter 5
Outlook
As for the future, mesospheric particles will remain in the focus of research
in the middle atmosphere. The research with sounding rockets will address
the small-scale interactions and processes needed in order to understand the
atmosphere as a whole. More needs to be learned about the particles
microphysics and interactions. This is particularly important since
mesospheric particle phenomena are considered to be a sensitive indicator
for the variability and long-term trends of the middle atmosphere. In order
to understand their role, the size spectrum and composition of both ice and
smoke particles needs to be quantified. Further development of direct
sampling techniques will hopefully open new ways to study these particles.
It is very important to be able to relate the measured particles to the actual
particle population in the mesosphere. There are a number of improvements
that can be made to the particle motion models described and used in
papers I, II, and III. The Brownian motion description can be further
developed to include a more realistic treatment of the molecule/particle
collisions. Both inelastic collisions and non-spherical particles can be
considered. The mass loss due to heating and subsequent sublimation in the
shocked gas flow in front of a detector is negligible for smoke particles and
large ice particles. However, for smaller ice particles this becomes
important and must be included in the model if ice particles are to be traced.
Also neglected in the model at this point are effects of payload charging by
photons, ions, and particles, as well as the possibility that incident particle
trajectories are influenced by such a payload charging. Another important
development is to extend the model into 3 dimensions to be able to simulate
rocket payloads with angles of attack other than 0° or payloads that are
asymmetric. The Brownian motion model can be used to quantify the
43
possible contamination of the Hygrosonde water measurement [Khaplanov
et al., 1996; Lossow et al., 2008a] due to adsorption or outgassing from the
rocket payload by studying how water vapour diffuses in the
aerodynamically shocked region in front of the instrument.
For a better understanding of mesospheric ice particle phenomena it is
important to develop new optical techniques to efficiently access the
smaller particles in the ice particle population. In paper IV we show how
this can be done by extending scattering experiments to be performed in the
Mie scattering regime, using wavelengths of the order of the particle size.
Due to a fortunate coincidence, the wavelength of the strong solar Lyman-α
line at 121.57 nm coincides with a minimum in the absorption spectrum of
O2 that otherwise efficiently shields the atmosphere from UV radiation in
the wavelength range 100-180 nm. As a result, direct Lyman-α penetrates
deep into the atmosphere during the day, well below the ice clouds in the
mesosphere. Further development of the SLAM instrument and
incorporation on sounding rocket payloads together with longer wavelength
measurements would give new and important information about the polar
mesospheric ice layer.
The new concept presented in paper V of calibrating direct atomic oxygen
measurements with simultaneous photometric O2 airglow measurements is
promising. In order to further improve the accuracy of this method, it is
important to further investigate the rate constants and excitation parameters
related to the airglow emissions. The coefficients currently in use are to a
large extent based on data from the single rocket campaign ETON. A
future, independent sounding rocket project with dedicated instruments to
study both atomic oxygen and oxygen related airglow is highly desirable.
For future rocket projects, the development of small and “cost-effective”
instrumentation will be mandatory. The concepts described in this thesis are
currently part of the preparation for the PHOCUS rocket campaign from
Esrange in 2010. They are also part of the continuously ongoing process of
providing improved experimental tools for the scientific understanding of
our atmosphere.
44
Acknowledgements
When I first started my Ph. D. position at MISU, five years felt like a very
long time. However, as you probably know, time flies when you do fun and
interesting things. These past five years has not been an exception and the
time has gone by very fast, maybe even too fast. I thank everyone at MISU
for making this part of my life very pleasant and memorable.
I wish to express my deepest and most sincere appreciation towards my
supervisor Jörg Gumbel for the inspiration and guidance through the
different projects and for being a very good friend. Many thanks also to my
assistant supervisor Markus Rapp at the Leibniz Institute of Atmospheric
Physics in Kühlungsborn, Germany. I am deeply grateful for being part of
the AP group together with Jörg Gumbel, Jacek Stegman, Misha
Khaplanov, Georg Witt, Bodil Karlsson, Stefan Lossow, Linda Megner,
Farah Khosrawi, Heiner Körnich, Tomas Waldemarsson and Peggy
Achtert. You have all certainly provided a wonderful and stimulating
working atmosphere. I have very much enjoyed all our travels, campaigns
and dinners (and of course work) together. I have learned so much from
you all, especially Jörg, Misha, Jacek and Georg. Thank you for sharing
your deep knowledge!
Special thanks also to Leif Bäcklin and Nils Walberg for the help in the
machine shop, to Eva Tiberg for knowing and taking care of almost
everything, to Birgitta Björling for help with finding old publications, to
Solveig Larsson, Pat Johansson, Susanne Eriksson, and Janne Söderlund for
help with a variety of things, and to Mikael Burger for the computer
support. You are all really a very important part of the department.
I wish to thank everyone at Andøya Rocket Range, Norway, and Swedish
Space Corporation at Esrange and in Solna for making my visits there
45
during meetings, experimental tests and campaigns very memorable. Many
thanks also to Karl Heinrich Fricke for the instructive time at Esrange.
I thank the Swedish National Graduate School of Space Technology for the
support in funding my Ph. D. studies.
And of course I thank Jonathan Friedman for introducing me to the middle
atmosphere during my M. Sc. thesis work at the National Astronomy and
Ionosphere Center, Arecibo Observatory, Puerto Rico. This certainly
inspired me to continue in the field of atmospheric science.
Last but not least, I thank my family and friends for supporting and
believing in me, and of course my Jenny, the worlds greatest, for
everything.
46
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