Space Science Reviews
Enceladus as an active body
--Manuscript Draft-Manuscript Number:
Full Title:
Enceladus as an active body
Article Type:
Regular review paper
Keywords:
Enceladus; Satellites; Saturn System
Corresponding Author:
Juergen Schmidt
University of Oulu
Oulu, FINLAND
Corresponding Author Secondary
Information:
Corresponding Author's Institution:
University of Oulu
Corresponding Author's Secondary
Institution:
First Author:
Sascha Kempf
First Author Secondary Information:
Order of Authors:
Sascha Kempf
Juergen Schmidt
T. Brockwell
Thomas Cravens
Larry W. Esposito
Katherina Fiege
Bernde Giese
Matthew M. Hedman
Paul Helfenstein
Terry A. Hurford
Ralf Jaumann
Geraint Jones
Sheila Kanani
Susan W. Kieffer
William Lewis
Brian A. Magee
Dennis L. Matson
Francis Nimmo
Robert Pappalardo
Mark E. Perry
Frank Postberg
Joachim Saur
John Spencer
Christophe Sotin
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Frank Spahn
Benjamin Teolis
Grabriel Tobier
Hunter Waite
David T. Young
Mikhail Zolotov
Order of Authors Secondary Information:
Abstract:
Data from the Cassini-Huygens mission to Saturn revealed that the mid-sized icy
satellite Enceladus is geologically active: A region around the south pole of Enceladus
is anomalously warm and a mixture of gases, dominantly water vapor, is found to flow
out there from a system of cracks in the ice crust. Micron sized ice particles, entrained
in the gas, are ejected to space. It became clear, that Enceladus is the dominant
source for magnetospheric plasma in the Saturn system, as well as for neutrals and
dust, which forms the E ring of Saturn. What is so particular about Enceladus? Why is
the activity localized at the south pole? How is the heat ultimately generated? How old
is the phenomenon and how did it evolve? A complete, consistent answer to these
questions is not known to date. Nevertheless, much progress has been made in
analyzing and understanding various aspects of the activity, which gives important
constraints and directions for future research. It is the purpose of this paper to review
the current understanding of Enceladus' geodynamics, geochemistry, and the
interaction with the saturnian system.
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Enceladus as an active body
Sascha Kempf · Jürgen Schmidt · T.
Brockwell · Thomas Cravens · Larry W.
Esposito · Katherina Fiege · Bernd Giese ·
Matthew M. Hedman · Paul Helfenstein ·
Terry A. Hurford · Ralf Jaumann · Geraint
Jones · Sheila Kanani · Susan W. Kieffer ·
William Lewis · Brian A. Magee · Dennis
L. Matson · Francis Nimmo · Robert
Pappalardo · Mark E. Perry · Frank
Postberg · Joachim Saur · John Spencer ·
Christophe Sotin · Frank Spahn · Benjamin
D. Teolis · Gabriel Tobie · Hunter Waite ·
David T. Young · Mikhail Zolotov
the date of receipt and acceptance should be inserted later
The first two authors organized the workshop Active Enceladus at the International Space
Science Institute (ISSI) in Bern, Switzerland, from which this paper results. All others are
listed alphabetically.
Sascha Kempf
LASP, University of Colorado, 1234 Innovation Drive, Boulder, CO 80 309-0392, USA
Jürgen Schmidt
University of Oulu, PL 3000, 90014 Oulu, Finland
T. Brockwell
Space Science and Engineering Division, Southwest Research Institute, San Antonio, TX 78238,
USA
Thomas Cravens
University of Kansas, Lawrence, Kansas, USA
Larry W. Esposito
LASP, University of Colorado, 1234 Innovation Drive, Boulder, CO 80 303-7814, USA
Katherina Fiege
IGEP, Universität Braunschweig, Mendelsohnstr. 3, 38 106 Braunschweig, Germany
Bernd Giese
DLR, Institut für Planetenforschung, Planetengeologie, Rutherfordstraße 2, 12 489 Berlin, Germany
Matthew M. Hedman
Cornell University, 322 Space Sciences Building, Ithaca, NY 14853-6801, USA
Paul Helfenstein
Center for Radiophysics and Space Research, Cornell University, Ithaca, NY 14853-6801, USA
Terry A. Hurford
Planetary Systems Laboratory, NASA Goddard Space Flight Center, Greenbelt, MD 20771,
USA
Ralf Jaumann
DLR, Institut für Planetenforschung, Planetengeologie, Rutherfordstraße 2, 12 489 Berlin, Germany
Geraint Jones
MSSL, University College London, Holmbury St Mary, Dorking, Surrey, RH5 6NT, UK
Sheila Kanani
MSSL, University College London, Holmbury St Mary, Dorking, Surrey, RH5 6NT, UK
Susan W. Kieffer
University of Illinois at Urbana-Champaign, 1301 W. Green St., Urbana, IL 61 801, USA
William Lewis
Space Science and Engineering Division, Southwest Research Institute, San Antonio, TX 78238,
USA
2
Sascha Kempf et al.
Abstract Data from the Cassini-Huygens mission to Saturn revealed that the mid-
sized icy satellite Enceladus is geologically active: A region around the south pole
of Enceladus is anomalously warm and a mixture of gases, dominantly water vapor,
is found to flow out there from a system of cracks in the ice crust. Micron sized ice
particles, entrained in the gas, are ejected to space. It became clear, that Enceladus
is the dominant source for magnetospheric plasma in the Saturn system, as well
as for neutrals and dust, which forms the E ring of Saturn. What is so particular
about Enceladus? Why is the activity localized at the south pole? How is the
heat ultimately generated? How old is the phenomenon and how did it evolve? A
complete, consistent answer to these questions is not known to date. Nevertheless,
much progress has been made in analyzing and understanding various aspects of
the activity, which gives important constraints and directions for future research.
It is the purpose of this paper to review the current understanding of Enceladus’
geodynamics, geochemistry, and the interaction with the saturnian system.
Brian A. Magee
Space Science and Engineering Division, Southwest Research Institute, San Antonio, TX 78238,
USA
Dennis L. Matson
Jet Propulsion Laboratory, 4800 Oak Grove Drive, Pasadena, CA 91109, USA
Francis Nimmo
Dept. of Earth and Planetary Sciences, University of California Santa Cruz, Santa Cruz, CA
95064, USA
Robert Pappalardo
Jet Propulsion Laboratory, 4800 Oak Grove Drive, Pasadena, CA 91109, USA
Mark E. Perry
Johns Hopkins University, Applied Physics Laboratory, Laurel, Maryland, USA
Frank Postberg
Mineralogisches Institut, Universität Heidelberg, Im Neuenheimer Feld 236, 69 120 Heidelberg,
Germany
Joachim Saur
Institute of Geophysics and Meteorology, University of Cologne, Albertus-Magnus-Platz, 50 923
Köln, Germany
John Spencer
Southwest Research Institute, 1050 Walnut St., Suite 300, Boulder, CO 80 302, USA
Christophe Sotin
Jet Propulsion Laboratory, 4800 Oak Grove Drive, Pasadena, CA 91109, USA
Frank Spahn
Universität Potsdam, Karl-Liebknecht Stras̈e 24/25, 14 469 Potsdam Golm, Germany
Benjamin D. Teolis
Space Science and Engineering Division, Southwest Research Institute, San Antonio, TX 78238,
USA
Gabriel Tobie
LPNG, CNRS, Universite de Nantes, UMR 6112, France
Hunter Waite
Space Science and Engineering Division, Southwest Research Institute, San Antonio, TX 78238,
USA
David T. Young
Space Science and Engineering Division, Southwest Research Institute, San Antonio, TX 78238,
USA
Mikhail Zolotov
School of Earth and Space Exploration, Arizona State University, Tempe, AZ 85287-1404, USA
Enceladus as an active body
3
Contents
1
2
3
4
5
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Geodynamics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.1 Observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.1.1 Global shape and topography . . . . . . . . . . . . . . . . . . . . . . . .
2.1.2 Thermal emission . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.1.3 Surface physical properties . . . . . . . . . . . . . . . . . . . . . . . . .
2.1.4 Tectonic structures and stratigraphic map . . . . . . . . . . . . . . . . .
2.1.5 Crater distribution, crater morphology . . . . . . . . . . . . . . . . . . .
2.2 Present day state . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.2.1 Tectonic structures, tidally driven activities . . . . . . . . . . . . . . . .
2.2.2 Possible internal structures and thermal state . . . . . . . . . . . . . . .
2.2.3 Heat generation by tidal friction and orbital equilibrium . . . . . . . . .
2.2.4 Heat transfer from deep interiors to the surface . . . . . . . . . . . . . .
2.3 Evolution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.3.1 Geological history . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.3.2 Internal dynamics and reorientation . . . . . . . . . . . . . . . . . . . .
2.3.3 Ocean lifetimes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.3.4 Coupled thermal-orbital evolution and episodic activities . . . . . . . .
Geochemstry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.1 Chemical observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.1.1 Surface composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.1.2 Composition of plume gases . . . . . . . . . . . . . . . . . . . . . . . . .
3.1.3 E ring and plume particles . . . . . . . . . . . . . . . . . . . . . . . . .
3.2 Composition of the interior . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.2.1 Bulk composition and a rocky core . . . . . . . . . . . . . . . . . . . .
3.2.2 Icy shell . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.2.3 Possible liquid phases . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.2.4 Aqueous vs. dry sources of plume emissions . . . . . . . . . . . . . . . .
3.3 Chemical evolution scenarios . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3.1 Accreted materials and chemical effects of water-rock differentiation . .
3.3.2 Chemical processes related to tidal heating . . . . . . . . . . . . . . . .
3.3.3 Habitability of Enceladus . . . . . . . . . . . . . . . . . . . . . . . . . .
3.4 Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Plume formation and properties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.1 The gas and dust components of the plume . . . . . . . . . . . . . . . . . . . .
4.2 Constraints . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2.1 Steadiness and variability . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2.2 Dust to gas mass ratio . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2.3 Grain size and velocity distribution . . . . . . . . . . . . . . . . . . . . .
4.2.4 Supply rate of particles to the E ring . . . . . . . . . . . . . . . . . . . .
4.2.5 Composition: Sodium and volatiles . . . . . . . . . . . . . . . . . . . . .
4.3 Scenarios for plume formation . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.3.1 Explosive boiling model . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.3.2 The clathrate reservoir hypothesis . . . . . . . . . . . . . . . . . . . . .
4.3.3 Shear heating models and sublimation from warm ice . . . . . . . . . .
4.3.4 Evaporation from liquid . . . . . . . . . . . . . . . . . . . . . . . . . . .
Interaction with Saturnian system . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.1.1 Discovery of Enceladus’s magnetospheric interaction . . . . . . . . . . .
5.1.2 Cassini’s encounters with Enceladus . . . . . . . . . . . . . . . . . . . .
5.2 The Plume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.2.1 Production rate and composition . . . . . . . . . . . . . . . . . . . . . .
5.2.2 Ionization in the Torus and the Plume . . . . . . . . . . . . . . . . . . .
5.2.3 Charged dust grains . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.3 Electrodynamic coupling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.3.1 Observations, models and particular aspects of Enceladus plasma interaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.4 E ring . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.5 Inner magnetosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.5.1 Energetic particle interaction . . . . . . . . . . . . . . . . . . . . . . . .
5.5.2 Neutral and plasma tori . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.5.3 Low energy electron enhancements . . . . . . . . . . . . . . . . . . . . .
5.6 Outer magnetosphere and solar wind . . . . . . . . . . . . . . . . . . . . . . . .
5.6.1 Oxygen at Titan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
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Sascha Kempf et al.
1 Introduction
The full extent of Enceladus’s activity was ultimately discovered in Cassini data.
But already much earlier its remarkable surface properties, seen in Voyager images,
led to the conclusion on geophysical activity (Squyres et al 1983; Haff et al 1983),
with the compelling suggestion that cryo-volcanic eruptions on Enceladus are the
source of the dusty E ring (Pang et al 1984; Showalter et al 1991a). That ring was
discovered by Feibelman (1967b) in images taken from earth, and earth-bound
observations at edge-on configurations showed that the ring is densest around the
orbit of Enceladus (Baum et al 1981), while the flybys of the Pioneer and Voyager
spacecraft had confirmed the close relation of the E ring and Enceladus.
In a detailed analysis of Voyager images Kargel and Pozio (1996) studied the
surface of Enceladus, assigning probable ages to various surface units from the
crater records (see also Kargel (2006)). They found very old terrains, that might be
nearly as old as the Saturn system, and regions practically void of craters, likely not
older than 107 − 108 years. They concluded that Enceladus must have undergone
multiple episodes of activity over a time span of billions of years, leading to a large
scale resurfacing, requiring large amounts of tidal heating, probably exceeding the
current steady state value. Remarkably, Showalter et al (1991a) interpreted the
narrow size distribution of the E ring, inferred photometrically as a strong support
for the hypothesis of geyser like cryo-volcanic ejection of ice grains.
In Cassini data the first indication for Enceladus’ geophysical activity was
found in perturbations of Saturn’s magnetic field near the moon. The Cassini
magenetometer team (Dougherty 2005) interpreted this observation in terms of
a partial atmosphere of Enceladus in the southern hemisphere, decelerating and
deflecting the corotational plasma, leading to the observed perturbation of the
planetary magnetic field.
In the subsequent targeted flyby E2 (Tab. 1) the intense geophysical activity
became evident in the data returned by the Cassini instruments, including the
discovery of the hot south polar region (Spencer et al 2006), associated with the
system of cracks dubbed ‘Tiger Stripes’ (Porco et al 2006), the peculiar diversity
of surface properties (Porco et al 2006) and its composition (Brown et al 2006),
the vapor plume emerging from the Tiger Stripes (Waite et al 2006; Hansen et al
2006), the dust plume (Spahn et al 2006b; Porco et al 2006) formed by micron
sized water ice grains, and the plumes’ interaction with the Kronian magnetosphere
(Dougherty et al 2006; Kivelson 2006; Tokar et al 2006; Jones et al 2006) and the
E ring (Spahn et al 2006b).
In the following years Enceladus attracted the attention of researchers from
a large variety of scientific disciplines. This paper gives a summary of progress
achieved and it is organized as follows. In Section 2 we cover various aspects of
Enceladus’s geodynamics, as its shape, tectonic structure and craters, internal
structure, tidal heating and heat transfer, internal dynamics and ocean evolution,
as well as the coupled orbital and thermal evolution with the possibility of a
recurrent transient activity. In Section 3 we summarize the constraints on the
geochemistry, composition of rocky core and ice shell, indications on aqueous and
dry processes, as well as plausible scenarios of chemical evolution. Section 4 is
devoted to the Enceladus plume, constraints on its gas and dust components, and
a discussion of scenarios for plume formation. Finally, in Section 5 the effect on the
Saturnian system is discussed, the replenishment of Kronian plasma, the neutral
torus, and the dusty E ring by material from Enceladus, the interaction with the
magnetosphere, and delivery of species to the outer magnetosphere.
Enceladus as an active body
Encounter
Orbit
E0
E1
E2
E3
E4
E5
E6
E7
E8
E9
E10
E11
E12
E13
E14
E15
E16
E17
E18
E19
E20
E21
E22
3
4
11
61
80
88
91
120
121
130
131
136
141
142
154
155
156
163
164
165
223
224
228
5
Date
2005-02-17
2005-03-09
2005-07-14
2008-03-12
2008-08-11
2008-10-09
2008-10-31
2009-11-02
2009-11-21
2010-04-28
2010-05-18
2010-08-13
2010-11-30
2010-12-21
2011-10-01
2011-10-19
2011-11-06
2012-03-27
2012-04-14
2012-05-02
2015-10-14
2015-10-28
2015-12-19
(048)
(068)
(195)
(072)
(224)
(283)
(305)
(306)
(325)
(118)
(138)
(225)
(334)
(355)
(274)
(292)
(310)
(087)
(105)
(123)
(287)
(301)
(353)
Time
(UT)
03:30
09:08
19:55
19:06
21:06
19:06
17:14
07:42
02:10
00:10
06:05
22:31
11:54
01:08
13:52
09:12
04:59
18:30
14:02
09:31
10:41
15:23
17:49
Closest Approach
Altitude
Speed
(km)
( km s−1 )
1260
6.6
497
6.6
166
8.2
48
14.4
49
17.7
25
17.7
169
17.7
99
7.7
1697
7.7
100
6.5
437
6.5
2555
6.8
46
6.3
48
6.2
99
7.4
1231
7.4
497
7.4
74
7.5
74
7.5
73
7.5
1839
8.5
49
8.5
4999
9.5
Geometry
U
U
U
P
P
P
P
S
S
S
U
S
N
N
Plume
Distance
( km)
330
242
193
167
310
100
1601
49
S
S
S
N
S
S
Table 1 Parameters of the Cassini orbiter encounters with Enceladus. The last three flybys
E20, E21, and E22 are listed here as planned for the Cassini Solstice Mission as of this writing.
The plume source is here assumed as the south pole of Enceladus’ surface. The flyby geometries
are defined as follows: U = upstream; P = North- South plume; S = southern; N = northern.
51
2 Geodynamics
52
2.1 Observations
53
2.1.1 Global shape and topography
54
55
56
57
58
59
60
61
62
63
64
65
66
67
68
69
70
71
72
73
74
75
76
77
78
79
The topography of Enceladus is not properly known yet. Topography requires an
equipotential surface for referencing heights, but elevation models of Enceladus
presented so far (Schenk and McKinnon 2009; Giese and The Cassini Imaging
Team 2010) refer to the best-fit ellipsoid (256.6 × 251.4 × 248.3 km) which, however,
is non-hydrostatic (Thomas et al 2007) and thus does not form an equipotential
surface. As a consequence, large-scale features may actually be smoother when
referenced to the true equipotential surface (which is currently unknown). Schenk
and McKinnon (2009) have studied whether a negative core density anomaly could
explain the observed 100 km wavelength depressions simply by deflection of the
equipotential surface, and concluded that this explanation did not work. The
depressions are more likely due to ice shell density and/or thickness variations
(Schenk and McKinnon 2009). However, it is unclear how a convecting ice shell
could sustain such lateral variations on the long term, and indeed the ice shell
might well be static (conductive) in most places (see section 2.3.2). Further modeling is needed here.
Currently existing elevation models (Schenk and McKinnon 2009; Giese and
The Cassini Imaging Team 2010) cover about 60% of Enceladus’ surface. They
show several up to 2 km deep depressions comparable in size to Enceladus’ radius.
The most pronounced one (centered at 10◦ S/230◦ E) is ovoid-shaped (200 × 140
km) and comprises both older cratered terrains and younger, resurfaced terrains
implying that it does not correlate with geologic boundaries (Schenk and McKinnon 2009). Likewise, there is a more circular 250 km depression (centered at
10◦ N/309◦ E) with a distinct lower-lying area (70 × 120 km) not correlating with
geologic boundaries. The South Polar Terrains (latitudes > 65o ) correlate with a
∼ 700 deep depression (Thomas et al 2007). There also exist large-scale convexshaped features on Enceladus (71◦ N/38◦ E and 30◦ S/300◦ E (Giese and The Cassini
6
Sascha Kempf et al.
88
Imaging Team 2010)). These 200 km bulges are about 1 km high and, contrary
to the depressions, correlate with resurfacing features. Local, small-scale features
on Enceladus reach elevation similar to those at large scales: (i) 15-20 km craters
reach depths of up to 1.8 km, (ii) the central peak of Aladdin crater (63◦ N/341◦ E,
D∼3 km) stands about 1 km above the local surroundings, (iii) the Harran Sulci
(3◦ N, 130◦ E), an 11 km wide half-graben structure, has an elevated flanking ridge
up to 1 km high (Giese et al 2008), (vi) an active 2 km wide segment in Enceladus’
South Polar Terrains forms a V-shaped trough up to 500 m deep (Giese and The
Cassini Imaging Team 2010).
89
2.1.2 Thermal emission
80
81
82
83
84
85
86
87
106
The total endogenic power of Enceladus’ south polar terrain has been recently
reevaluated using the 2008 Cassini Composite Infrared Spectrometer (CIRS) 10
to 600 cm−1 spectra (Howett et al 2011). The new estimate of 15.8 ± 3.1 GW is
about three times higher than the previous estimate of 5.8 ± 1.9 GW (Spencer
et al 2006), which was derived from > 600 cm−1 observations, because the new
estimate includes low-temperature emission that the earlier higher-wavenumber
observations were not sensitive to. This indicates that Enceladus’s south polar
terrain is even more thermally active than initially thought.
The latest CIRS > 600 cm−1 observations confirm that the maximum of emission occurs along the tiger stripes and most of the emissions are confined to a narrow, intense region no more than a kilometer (half a mile) wide along the faults.
Emission is continuous along the tiger stripes at 5 -10 km resolution (Howett et al
2011), but shows large spatial variations along-strike at higher resolution (Spencer
et al 2011). The peak temperatures observed by CIRS along Damascus and Baghdad Sulci, the most active tiger stripes, are estimated to about 190 K, over widths
if tens of meters along the stripes, but higher temperatures are possible over narrower regions closer to the fractures.
107
2.1.3 Surface physical properties
90
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93
94
95
96
97
98
99
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Enceladus’ surface is mostly composed of pure water ice, amorphous and crystalline, except near its south pole where light organics and CO2 have also been
identified, particularly in the Damascus, Baghdad, Cairo and Alexandria Sulcus
regions (Brown et al 2006). Variations in the main water ice absorptions at 1.04,
1.25, 1.5, and 2.0 µm indicate that the averaged particle size varies on Enceladus’
surface and is directly correlated with geological units (Jaumann et al 2008). On
a regional scale heavily cratered terrains exhibit a mean particle size of 15 ± 5 µm,
larger particles with diameters up to 30 ± 10 µm are correlated either with fresh
impact crater ejecta or with ridges and fractures marking the transition zone to
the lightly cratered plains. Tectonically deformed regions exhibit particles ranging
from 40 µm about 120 µm with the largest ones (about 2% reaching values of about
0.2 mm) located close to the center of the south polar sulci (Brown et al 2006;
Jaumann et al 2008; Spencer et al 2009).
Similar patterns are seen in the ISS multispectral images of the south polar region, which show enhanced near-infrared absorption near the tiger stripes
and other tectonically disrupted regions, consistent with larger grain sizes (up
to100 µm) near the sulci (Porco et al 2006). Possible causes for the large grain
sizes near the south polar sulci include preferential fallout of larger, slower, plume
grains near their sources (Jaumann et al 2008; Hedman et al 2009), or sintering
and growth of grains resulting from deposition of plume gases or from the enhanced temperatures near the south polar tectonic features (Spencer et al 2009).
On a global scale, the visible orbital phase curve of Enceladus shows that the
trailing hemisphere is brighter than the leading hemisphere (Buratti and Veverka
1984; Buratti et al 1998), and ground-based spectra show smaller ice particles
on the trailing hemisphere (Verbiscer et al 2006). These variations may reflect
hemispheric differences in Enceladus’ interaction with the E-ring.
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Disk-integrated spectra of Enceladus indicate that the water ice in the equatorial and mid- to high-latitude regions is primarily crystalline (Grundy et al 1999;
Emery et al 2005). However, Cassini spectra has detected locally abundant amorphous ice in the south polar region between the sulci, while the sulci themselves
are dominated by crystalline ice (Brown et al 2006; Newman et al 2008). Amorphous water ice forms when it is condensed directly from the vapour to a solid at
temperatures below about 100K (Brown et al 2006; Newman et al 2008). Away
from the sulci, the surface temperatures are well below 100 K (Spencer et al 2006),
so amorphous ice should be able to form and remain stable for long periods of
time there.
The icy surface of Enceladus is affected by various modification processes, including charged particle bombardment (implantation or sputtering), E-ring grain
bombardment or coating, thermal processing, UV photolysis, and micrometeoroid
bombardment. Water ions are the dominant sputtering agents, and rates are estimated to be about 0.1 µm per year, or 10 cm per million years (Johnson et al
2008). The high albedo (Squyres et al 1983; Verbiscer et al 2005; Spencer et al
2006) is probably due to a coating of plume particles, derived both directly from
plume fallout and from re-impact of plume particles stored temporarily in the
E-ring (Spahn et al 2006b; Kempf et al 2008; Spencer et al 2009; Jaumann et al
2009), suggesting a domination of coating over micrometeorite gardening. On the
other hand, IR-spectra reveal a consistent pattern of increasing grain size mostly at
tectonic features with decreasing surface age (Jaumann et al 2008), which suggest
a gradual comminution of particles over time by impact gardening or sputtering.
This might be due to either uneven coating at steeper slopes or mass wasting in
those areas.
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2.1.4 Tectonic structures and stratigraphic map
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Enceladus’ surface reveals a wide variety of icy tectonic structures that record a
long history of tectonic deformation. These structures include rifts and tension
cracks (fossae), lanes of striated and grooved terrains (sulci), networks of narrow
ridges and domelike structures (dorsa), icy massifs and mountain chains, and peculiar quasi-parallel systems of parallel ridges that are most often associated with
active south polar tectonism and volcanism. Slicing through nearly all of these
features are ubiquitous sub-parallel arrangements of gossamer cracks with typical
spacing of hundreds of meters (Miller et al 2007; Barnash et al 2006). Spencer et al
(2009) provided a detailed review of global tectonic structures that were known
at the time. Since then, new Cassini imaging coverage has revealed previously unknown details of leading hemisphere and South Polar terrains, and new progress
has been made in tectonic mapping and modeling of already-recognized features.
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Detailed stratigraphic mapping (Fig. 1) of the South Polar Terrain region has
been undertaken by Patthoff and Kattenhorn (2009, 2010, 2011). They distinguished between present day active tiger stripes and three earlier systems of
paleo-tiger stripes, each system of which share a common orientation and relative age. The old tiger stripes are crosscut and offset by progressively younger
fracture sets. Incremental changes in orientation of these systems suggest that
Enceladus lithosphere has experienced non-synchronous rotation over time. The
differing orientations of the modern and paleo-tiger stripe systems, respectively
were found to correlate with corresponding fracture orientations in the deformation belts that separate the SPT province from the older terrains to the north.
These correlations are consistent with hypothesis that tectonic spreading related
to tiger stripes (Helfenstein et al 2008a; Spencer et al 2009; McKinnon and Schenk
2009; Barr 2008; Barr and Preuss 2010) has occurred over time and that corresponding convergence is manifested as fold-like belts and complex fracture systems
at the boundary of the SPT (Helfenstein et al 2006; Porco et al 2006; Schenk and
McKinnon 2009).
The most prominent ancient tectonic features on Enceladus are likely found
within tectonized plains on the leading hemisphere and the trailing hemisphere.
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Fig. 1 South Polar fractures and their relation to geological history and the dichotomous
boundary of the STP region from Patthoff and Kattenhorn (2011). Dark red lines are the
four active tiger stripes while pink lines are fractures which parallel them. Yellow, green, and
blue lines, respectively, identify stratigraphic fracture sets 2, 3, and 4. The dark purple lines
delineate the SPT structure dichotomy that separates the inner fractured terrain from the
outer, older terrain. Fold-like features are shown in light green and fractures that crosscut
them are represented by light purple. Tiger stripe systems of different stratigraphic age and
orientation correlate with features of the SPT boundary dichotomy, suggesting that the same
tectonic mechanisms that operate at present have operated in the past when the tiger stripe
systems were oriented in different directions, perhaps due to non-synchronous rotation.
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Tectonic overprinting and evidence for early viscous relaxation is widespread and
often only partial relict features may be recognized. As noted earlier, the broad
hemisphere-centered regions of tectonic resurfacing are subdivided into smaller
quasi-polygonal provinces that exhibit distinct tectonic styles and differing ages.
On the leading hemisphere is a conspicuous province near 30◦ N, 90◦ W of curvilinear massifs and roughly orthogonal-trending wide, shallow troughs (cf. Pappalardo
and Crow-Willard 2010) that define a crudely radial and concentric pattern relative to a point near 25◦ N, 125◦ W. This angular sector, about 65◦ in width, may
be either the partial remains of an ancient impact basin with a diameter of about
180 km or perhaps a surface expression of an ancient, large diapir. These hypotheses are also consistent with the presence on the leading hemisphere of numerous
subdued-relief circular albedo patterns and smooth circular patches up to 40 km in
diameter (cf. Spencer et al 2009, Fig. 21.7). The peculiar quasi-radial wide-shallow
troughs resemble extinct, topographically degraded examples of tiger stripes seen
near the South Pole. While these features may have a different fracture origin from
tiger stripes, their comparable morphology suggests that long ago they may have
expressed a similar style of fissure volcanism (Helfenstein et al 2010b,a).
A peculiar system of rounded, ropy ridges separate the tiger stripe fractures in
the SPT region. They are curved, quasi-parallel knobby ridges and clustered arrangements of elongated barrel-shaped knobs that were first recognized as grooved
bands by Porco et al (2006), but are called funiscular plains (Helfenstein et al
2008b,a; Spencer et al 2009; Helfenstein et al ????a) and South Polar folds (Barr
and Preuss 2010) in more recent works. They trend approximately parallel to
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tiger stripes and their lengths and kilometer-scale widths are comparable to those
nearby tiger stripe flanks. Near their distal ends where funiscular plains intersperse
appear to break into a more chaotic arrangement of elongated to equant knobs and
cylindrical segments. The superficial resemblance of offset patterns at the edges
of tiger stripes to terrestrial transform faults and stepovers, the morphological
similarity of an axial discontinuity along the Damascus Sulcus tiger stripe to terrestrial offset spreading centers and a plausible preliminary tectonic reconstruction
of South Polar paleo-terrains, suggest that the ropy ridges may form as folds due to
tectonic spreading around active tiger stripes, very crudely analogous to terrestrial
sea floor spreading (Helfenstein et al 2008a, ????b; Spencer et al 2009; McKinnon
and Schenk 2009). This interpretation is supported by the atypical regional topographic flatness of the active South Polar Terrain (McKinnon and Schenk 2009), as
well as models of mobile-lid convection (Barr 2008; O’Neill and Nimmo 2010) and
terrestrial lava flow folding (Barr and Preuss 2010). The latter found that icy folds
due to spreading could form in an ice shell with an upper high-viscosity boundary
layer of thickness < 400 meters, with a driving stress of 40 to 80 kPa, and strain
rate between 10−14 s−1 and 10−12 s−1 .Their results suggest that resurfacing of
the SPT could occur over 0.05 to 5 Myr, consistent with its estimated surface age
from cratering statistics. A stratigraphically much older region containing rounded
ropy patterns of ridges that are similar in morphology and scale to those in the
SPT have recently been found on the leading hemisphere of Enceladus, near 10◦ S,
60◦ W, suggesting that perhaps tectonic spreading may have occurred elsewhere on
Enceladus at an earlier era. However, the overall physical arrangement of tectonic
features on the leading hemisphere significantly differs from the distinct pattern
seen in the SPT, so that it is not likely to simply be an early clone of the SPT
system Pappalardo et al (2010); Helfenstein et al (2010a).
Dorsa are straight to curvilinear narrow ridges hundreds of meters wide, with
rounded sub-kilometer scale relief. They are found within all major tectonized
provinces, but the best studied examples are Cufa Dorsa (3.2◦ N, 286.2◦ W) and
Ebony Dorsum (5.7◦ N, 280.5◦ W). Their lengths generally do not exceed a few to
tens of kilometers. They occur as isolated features or in crudely polygonal arrays
along with elongated domelike features with similar height and width dimensions.
Prominent examples on Enceladus trailing hemisphere may have originated as
thrust faults (Pappalardo and Crow-Willard 2010). However, some dorsum variants, specially near the South Polar Terrain, appear to arise along pre-existing
fractures, suggesting that they were extruded (Porco et al 2006) or that they
may be transpressional features analogous to terrestrial positive flower structures
(Helfenstein et al 2008b,a, ????a). This is especially true of a special class of narrow
dorsa, called shark fins that are found medially within tiger stripe fractures.
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2.1.5 Crater distribution, crater morphology
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Crater counts on Enceladus confirm the wide range of surface ages (Plescia and
Boyce 1983; Lissauer et al 1988; Kargel and Pozio 1996; Porco et al 2006; Kirchoff
and Schenk 2009). Depending on the flux model used, which can be either lunarlike and therefore asteroidal (Neukum 1985) or, more plausibly, based on models of
cometary impact rates (Zahnle et al 2003), cratered plains materials have absolute
ages as old as 4.2 billion or 1.7 billion years, respectively. In general, the eastern
fractured terrain has formed either between 3.7 and 1.0 billion years ago in the
lunar-like scenario or it is as young as 200 to 10 million years based on the Zahnle
et al. model. The south polar terrain shows crater ages younger than 100 million
to few hundreds of thousands years, respectively, and consistent with the ongoing
activity at the tiger stripes.
Enceladus’ cratered plains distribution is unique in that it appears to have a
relative deficiency of craters for diameters ≤ 2 km and ≥ 6 km compared to the
other satellites heavily cratered plains (Kirchoff and Schenk 2009). The data reported by Kirchoff and Schenk (2009) also indicates that the impact crater density
within the cratered plains changes with latitude. Specifically, both the north and
south mid-latitude regions have approximately three times higher density than
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the equatorial region. The “missing” small and large craters in Enceladus cratered
plains may be due to a combination of viscous relaxation of the larger craters,
and burial of the relaxed large craters and small craters by south polar plume and
possibly E-ring material. A large number of craters appears to viscously relaxed
in the apparently young regions of the anti-Saturnian and trailing hemispheres, as
well as in the older, upper northern latitudes (Bray et al 2007; Smith et al 2007;
Smith 2008). The occurrence of viscous relaxation in various terrains suggest large
geographical and time variations in heat flow.
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2.2 Present day state
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2.2.1 Tectonic structures, tidally driven activities
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One prominent tiger stripe near Enceladus’ south pole consists of arcuate segments, resembling the shape of cycloidal cracks on Europa (c.f. Hurford et al
2007). On Europa, these distinctive tectonic patterns were likely produced by periodic tidal stresses (Hoppa et al 1999). The similarity between Enceladus tiger
stripes and Europa’s cycloidal ridges is in shape only; they are quite distinct from
one another morphologically. Nevertheless, the similar shape of the tiger stripe
suggests that its formation may have been similarly controlled by diurnal tidal
stresses (Hurford et al 2007; Nimmo et al 2007). If this is the case, diurnal tidal
stresses may play an important role, driving activity on Enceladus.
Diurnal tidal stress on Enceladus is a natural product of its orbital eccentricity.
Enceladus’ finite orbital eccentricity causes small daily changes in the distance between Enceladus and Saturn, affecting the height of the tide raised on the satellite
by the planet. Besides affecting the height of the tide on Enceladus, the orbital
eccentricity also causes the longitude of the tidal bulge to oscillate as it tracks
the position of Saturn throughout an orbit. This constant reshaping of Enceladus’
shape produces stress on the surface. However, the tidal deformation and hence
the resulting stress get significant only if a liquid water layer decouples the ice
shell from the rocky core (Hurford et al 2007; Nimmo et al 2007; Tobie et al 2008).
Olgin et al (2011) recently showed that tidal stress can reach values of 1 bar for ice
shell thinner than 20 km, and that the stress could exceed the Coulomb criterion
for shear failure along the tiger stripes if the ice shell thickness is less than 40 km.
However, although tidal stress may be sufficient to re-activate pre-existing faults,
it is probably insufficient to generate new faults.
Other source of stresses must be taken into account in order to explain the
fracturing of the upper lithosphere. Change of volume due to thermal expansion/contraction or melting/crystallization of the internal ocean can lead for instance to stresses exceeding several hundred of bars (Nimmo 2004; Manga and
Wang 2007; Mitri and Showman 2008b). Changes in the solid surface position with
respect to the tidal bulge can also result in significant stresses. Non-synchronous
rotation (NSR) involves longitudinal shifts in position, while reorientation of the
rotational axis, called ”true polar wander”, also includes latitudinal shifts. Both
mechanisms result in predictable global stress patterns (e.g. Melosh 1980; Leith
and McKinnon 1996). Such a reorientation of the rotational axis results from perturbation of the moment of inertia of the body owing to large-scale internal or
surface loads, and it is also thought to have a play role in the evolution of Enceladus (Nimmo and Pappalardo 2006).
Once the lithosphere is pre-fractured, tidal stress is likely to be the main driver
of the propagation and persistence of cracks, by periodically reactivating them. Although the amplitude of stress is relatively small, the energy associated with tidal
motions is quite large and contributes significantly to the dynamics of the lithosphere. The Tiger Stripes in the south polar region exhibit higher temperatures
than the surrounding terrain and are the sources of the observed eruptions (Porco
et al 2006; Spencer et al 2006; Spitale and Porco 2007). Over Enceladus’ orbital
period, stresses across the tiger stripes alternate from compressive to tensile, possibly allowing the faults to open, which could expose a subsurface volatile reservoir,
Enceladus as an active body
11
Fig. 2 Schematic interior structure of Enceladus. See text for more details. The behaviour
of the ice shell is surprisingly poorly understood. As discussed below, it is not clear whether
or not the shell is convecting at the present day. Additionally,although there is clearly a heat
source at the South Pole, there is currently no agreement as to where the source of the heat is
located.
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creating an eruption (Hurford et al 2007). In a complementary process, periodic
shear stress along the rifts may generate heat along their lengths (Nimmo et al
2007) (see section 2.2.3). Model calculations of tidal shear stress along rifts in the
south polar region of Enceladus shows that is a reasonable correlation between
observed hotspot locations and predictions of regions, which show high amounts
of tidal shear (Nimmo et al 2007; Smith-Konter and Pappalardo 2008).
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2.2.2 Possible internal structures and thermal state
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The bulk density of Enceladus suggests an interior with roughly equal masses of
ice and silicate (Schubert et al 2007). Because the shape of Enceladus is not close
to that expected for a body in hydrostatic equilibrium ((a−c)/(b−c) 6= 4, (Thomas
et al 2007), see Section 2.1.1), neither shape nor gravity measurements can be used
to infer the moment of inertia of the body. Nonethless, given the high present-day
heat flux and indications of subsurface liquid water (see below), a differentiated
structure is generally assumed (Schubert et al 2007; Spencer et al 2009) (Figure
2).
The strongest argument for the presence of liquid water beneath the surface of
Enceladus is the detection of salty grains in the E-ring (Postberg et al 2009) (see
Section 3.1.3). Liquid at any location in the ice shell would tend to drain downward,
forming a local sea or a global ocean. We view the former of these possibilities as
more likely (see Section 2.3.3). A hot silicate core has been argued based on the
need to melt the ice shell beneath the south pole (Collins and Goodman 2007) and
to explain the observed NH3 in the plumes (Matson et al 2007). Neither argument
is strong: melting within the ice shell is a viable alternative (Tobie et al 2008) and
other pathways to generating NH3 have been suggested (Glein et al 2008). The
core of Enceladus is so small that any primordial heat will have been long lost via
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!"
Sascha Kempf et al.
#"
$"
%&'()*+,-./012345"6
!
Fig. 3 Three examples of heat production by tidal friction: a) Dissipation of eccentricity tides
associated with strong tidal flow in an thin and shallow internal ocean (Tyler 2011), b) volumetrically-distributed viscous dissipation in a convective ice shell (after Běhounková et al 2010),
c) shear heating associatied with tidal motions along tiger stripes (Nimmo et al 2007)
!
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conduction, and it is very unlikely that the core is sufficiently ductile that tidal
heating is important (Roberts and Nimmo 2008; Tobie et al 2008).
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2.2.3 Heat generation by tidal friction and orbital equilibrium
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The high south polar heat fluxes suggest either correspondingly high present-day
heat production rates, or some mechanism of storing heat and then releasing it
rapidly. The issue of heat transport is addressed below (Section 2.2.4); here we
focus on heat production.
Three main mechanisms of heat production have been proposed. A first one
is dissipation associated with strong tidal flow in an internal ocean (Tyler 2009,
2011). Large dissipation might be obtained for a thick ocean (1-50 km) but required
a non-zero obliquity. The obliquity is currently unknown, and it is expected to be
rather small. Enceladus is likely to be in a Cassini state, where the obliquity should
be equal or smaller than the orbital inclination (≤ 0.01◦ ). For such a low obliquity
value, obliquity tides remains unimportant (Tyler 2009, 2011). Tyler (2011) also
predicts that large dissipation associated with eccentricity tidal forces might also
be obtained but only for very thin oceans (≤ 360m, Fig. 3a). However, as Tyler’s
model assume that the ice lithosphere is thin enough to not affect the ocean tidal
response, his calculations are therefore only valid for shallow oceans (few hundred
meters under the surface) which is incompatible with topographic evidence of
bulge and depression on the order of 1-2 km, as well as with our understanding of
Enceladus’ interior structure. Such a thin, shallow ocean would be above an less
dense ice layer, which is gravitationally unstable.
A second dissipation mechanism is conventional tidal heating, involving volumetrically-distributed viscous deformation in response to the time-varying tidal
potential (e.g. Tobie et al 2005; Mitri and Showman 2008a; Běhounková et al
2010). As this dissipation mechanism depends on the viscosity (or temperature),
it strongly varies within a convective ice shell and therefore has a strong impact
on the dynamics of thermal instabilities (Fig. 3b) (Běhounková et al 2010).
The third mechanism is localized, shear heating on faults, again driven by tidal
deformation (Fig. 3c) (Nimmo et al 2007). In this case, the dissipation process is
also associated with viscous friction, but the tidal deformation is controlled by
strike-slip motions along the faults rather than volumetrically-distributed viscous
deformation.
In the ice shell, the two latter processes are probably operating simultaneously. Like the ocean flow mechanism, these two solid-state mechanisms require
the presence of a subsurface ocean, at least in the southern hemisphere : an ice shell
Enceladus as an active body
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coupled directly to the silicate mantle does not deform sufficiently to generate the
observed heating rates. Both mechanisms have difficulty explaining the degree-one
pattern of heating, because tidal deformation is a degree-two feature i.e. hot spots
should be observed at both poles. The most likely explanation is probably some
kind of pre-existing thermal anomaly, which is not terribly satisfying.
Although the details of the process need further study, shear heating appears
capable of explaining the observed south polar heat and vapour fluxes for reasonable parameter choices (Nimmo et al 2007). Generating the required power in a
localized warm diapir, either with a global ocean or a regional sea, is also possible
as long as the viscosity is relatively low (∼ 1013 − 1014 Pa s or less) (Tobie et al
2008). Dissipation in the ocean might also generate several gigawatts, but it requires very specific conditions in term of obliquity or ocean thickness, which could
be met only during relatively short periods of time.
Irrespective of the mechanism of heat production, an extremely important
point is that the observed current heat loss greatly exceeds the steady-state tidal
heat production rate (Meyer and Wisdom 2007). This steady-state value arises
because of the competition between the Enceladus:Dione 2:1 resonance and dissipation in Saturn, which increases the eccentricity e of Enceladus, and dissipation
within Enceladus, which damps it. The steady state heat production rate depends
on dissipation within Saturn. For a dissipation factor estimated to 18 000 in Saturn
(e.g. Meyer and Wisdom 2007), Enceladus is currently losing heat at a rate faster
than the long-term rate at which it can be produced, estimated to be 1.1 GW;
further discussion of this issue is deferred to Section 2.3.4.
Physical libration and forced libration associated with secondary spin-orbit
resonance (Wisdom 2004; Hurford et al 2009) may also lead to enhanced tidal deformation and hence strong tidal heating. Hurford et al (2009) showed, for instance,
that for the case of a physical libration at the Cassini’s limit of L = 1.5◦ , the tidal
heating is potentially enhanced by ∼ 4.9 times. Current efforts to directly detect
Enceladus’ physical libration from Cassini imaging limits its amplitude to ≤ 15◦
(6.6 km) (Porco et al 2006). Rotational models of Enceladus with no global ocean
predict that the librations related to the perturbations of the orbit by Dione may
reach amplitude of the order of 1km (Rambaux et al 2010), significantly smaller
than the value assumed by Hurford et al (2009). There is no indication that enhanced dissipation associated with libration play a key role at present, but it is
possible that Enceladus passed through such a high dissipative stage in a recent
past, especially if it enters a secondary spin-orbit resonance (Wisdom 2004).
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2.2.4 Heat transfer from deep interiors to the surface
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The internal heat, mainly produced by tidal dissipation in the H2 O layer, may be
transferred to the surface from the source region by conduction through the ice, or
by solid-state convection of water ice, or by flow of liquid water or by circulation
of water vapor. All these heat transport mechanisms probably contribute to the
total observed thermal emission, but their relative contribution to the total heat
transport strongly depend on the thermal structure of the icy crust. Thermal conduction is the main heat transport in regions where no other transport mechanism
is operating (i.e. in cold lithospheres). It plays also a major role in regions where
thermal gradients are large, for instance in the vicinity of the thermal anomalies
associated with the tiger stripes (Abramov and Spencer 2009). However, the thermal conductivity of water ice is not large enough to efficiently remove the heat
generated within Enceladus, so that other more efficient transport mechanisms are
required.
The intense plume activity on Enceladus suggest that vapor production and
circulation through the ice is an efficient mechanism to transport heat to the
surface (Nimmo et al 2007) (Fig. 4a). However, such a transport mechanism can
efficiently operate only in regions with high porosity, therefore mainly in the upper
crust where the confining pressure is low. It is conceivable that vapor can be
produced at shallow depths (down to few hundred meters) , but it gets rather
unlikely at deeper levels. Indeed, water vapor can be produced from sublimation
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Sascha Kempf et al.
a) Heat transport by vapor and conduction
in the shear heating model (Nimmo et al 2007)
b) Principle of the heat pump model
(Matson et al. 2010)
c) Convective heat transport in a tidally heated ice shell (Behounkova et al. 2010)
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Fig. 4 Examples of heat transport mechanisms: a) Heat removal by vapor production and
advection as predicted by the shear heating model (Nimmo et al 2007), b) Principle of the
heat pump model as proposed by Matson et al (2010). c) Convective thermal instabilities
in a tidally heated ice shell modeled by a 3D spherical code solving simultaneously thermal
convection and viscous dissipation (Běhounková et al 2010).
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of water ice only if the vapor pressure remains at or below the saturation (or
equilibrium) vapor pressure. For T = 200 K, the saturing vapor pressure is only
0.16 Pa (e.g. Schorghofer 2007). As a comparison, the hydrostatic pressure at one
kilometer below the surface is of the order of 100 kPa. Maintaining vapor pressures
as low as 0.2 Pa within pores a few hundred meters and more below the surface
would require all pores to be highly interconnected, resulting in an extremely
fractured permeable medium. Pressures in active wide fractures could possibly
remain close to the near-vacuum surface conditions under the action of tensile
tidal stresses if vapor is instantaneously removed (Kieffer et al 2006; Hurford et al
2007). However, this is possible only locally in the first hundred meters below the
surface. Therefore, vapor production and transport can be considered as a major
factor in advection of heat only near the surface.
Below the porous and fractured upper crust, heat is more likely advected
through thermal convective instabilities of solid ice and possibly by circulation
of water coming from the underlying ocean. Solid-state convection in Enceladus’
ice shell is predicted to occur only if the viscosity of ice at the base of the ice shell is
less than 1014 Pa s, corresponding to rather small grain size (< 0.3 mm, Barr and
McKinnon 2007) if the deformation of ice is controlled by grain boundary sliding.
The efficiency of heat removal by solid-state convection depend not only on the
basal viscosity but also on the mobility of the upper layer (Barr 2008; Roberts and
Nimmo 2008; Tobie et al 2008; O’Neill and Nimmo 2010; Běhounková et al 2010)
(Fig. 4c). In the asymptotic stagnant lid regime, scaling laws for a volumetrically
heated fluid (Grasset and Parmentier 1998; Solomatov and Moresi 2000) predict
that the heat flux ranging between roughly 15 and 30 mW m−2 . This can explain
less than 20% of the heat flux observed at the SPT. Moreover, convection in the
Enceladus as an active body
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15
standard stagnant lid regime is unable to reproduce the lateral variations as observed in the thermal emission data (Spencer et al 2006; Abramov and Spencer
2009). Such a convection regime does not take into account the mobility of the
overlying lithosphere, which is likely to participate in the convective cycle (Barr
2008; Tobie et al 2008; O’Neill and Nimmo 2010). On Earth, for instance, the motions of lithospheric plates play a crucial role in the advection of heat (e.g. Sotin
and Parmentier 1989). Plate spreading at terrestrial midocean ridges is known to
influence the upwelling of warm mantle materials and hence increase the release
of internal heat (Sotin and Parmentier 1989). Although the tectonic processes on
Enceladus are probably different, active spreading can also be envisaged over Enceladus’ south polar region (McKinnon and Barr 2007). An average spreading rate
of several centimeters would be sufficient to explain the observed thermal emission at the south pole (McKinnon and Barr 2007). The spreading would also help
focus the stream lines below the ridges (Sotin and Parmentier 1989), leading to
a convergence of warm materials heated over a relatively broad dissipative region
toward the narrower tectonic active region at the south pole.
An alternative heat transport may be closed-loop circulation of water from a
subsurface ocean (Matson et al 2010) (Fig. 4b). Ocean water is denser than the
icy crust and normally cannot rise to the surface. Liquid water can, however, be
injected at the base of the ice layer under the action of tidal forces leading to the
formation of bottom crevasse (Mitri et al 2008) and owing to overpressurization
of a crystallizing internal ocean (Manga and Wang 2007). As water moves upward
owing to overpressurization, the pressure is reduced and gases initially dissolved
in water can exsolve, form bubbles, and make the ocean water less dense than
the surrounding ice (Crawford and Stevenson 1988). Buoyant ocean water could
therefore rise toward the surface and carries with it heat and chemicals that are
delivered to the plume chambers and the shallow reservoirs that keep the surface
ice warm. Depleted of heat and gas bubbles, the water becomes more dense and
descends through cracks that return it to the ocean. Once the buoyancy has been
sufficiently established the closed-loop circulation is self-sustaining. The heat delivery rate is controlled by the rate at which water arrives and by how much heat
it carries. A discharge of ∼2000 m3 s−1 of seawater would be required to deliver
the observed 15 GW (Howett et al 2011). This can be compared with the supply
needs of the plumes. The amount of mass ejected by the plumes has been estimated to be about 150 to 350 kg s−1 whereas the heat flow needs a much larger
mass of water, namely about a hundred tons per second. Another way of gauging the reasonableness of the supply requirement is to calculate how much of the
surface area must be allocated to the upwelling of ocean water. Assuming a flow
velocity of a centimeter per second, the actual area is relatively small, being only
several millionths of the area of the South Polar Region (that has a radius of about
50 km). Higher flow velocities require less area.
Except for thermal conduction, the three other heat transport mechanisms
require very specific conditions to operate. Solid-state convection in the ice shell
can be initiated only if the viscosity is lower than 1014 Pa.s, which is possible, but
require small grain sizes ( < 1 mm) and temperature near the melting in most
of the convecting zone. Maintaining small grain sizes in a convective ice shell is
problematic as normal grain growth is relatively rapid at elevated temperature
(Montagnat and Duval 2000; Barr and McKinnon 2007). Circulation of water is
also problematic because as mentionned above liquid water is naturally denser
than water ice. The injection of liquid water at the bottom of the ice shell requires
relatively large tidal stresses or rapid ocean crystallization, and the extraction of
liquid water up to the surface, or at least at shallow depths, therefore requires the
presence of dissolved gases in sufficient amounts (at least a few 0.1 wt%, Tobie
et al (2010)). Finally the vapor transport model can operate only in regions with
very high porosity, therefore it is limited to the first hundred meters below the
surface where the confining pressure is low. These different transport mechanisms
are probably operating simultaneously, but it is unclear which of these transport
mechanisms is the dominant one at present at Enceladus’ south pole, and how the
efficiency of heat transport evolves as a function of time.
16
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2.3 Evolution
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2.3.1 Geological history
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Sascha Kempf et al.
Enceladus is an active icy satellite operating geological processes that are also
thought to have occurred on other icy satellites. Thus, Enceladus is important to
better understand the histories of other icy satellites in the Solar System. Enceladus exhibits the greatest geological diversity in a complex stress and strain history that is predominantly tectonic and cryovolcanic in origin, with both processes
stratigraphically correlated in space and time, although the energy source is so far
not completely understood (e.g. Smith et al 1982; Kargel and Pozio 1996; Porco
et al 2006; Nimmo and Pappalardo 2006; Nimmo et al 2007; Jaumann et al 2008,
2009; Spencer et al 2009).
Three different terrain types exist on Enceladus. Ancient cratered terrains cover
a broad band from 0◦ to 180◦ longitude. Centred on 90◦ and 270◦ longitude there
are younger, deformed terrains roughly 90◦ wide at the equator. The southern
polar region south of 55◦ is heavily deformed and almost uncratered. The most
remarkable known aspect of Enceladus is its south polar plume activity (Hansen
et al 2006; Waite et al 2006; Porco et al 2006; Dougherty et al 2006; Spahn et al
2006b). This is the only unambiguous example of active cryovolcanism in the solar system. The plumes arise from warm surface fractures (Spencer et al 2006),
the ”tiger stripes” (Porco et al 2006). Sublimation of warm surface ice (Spencer
et al 2006), boiling of near-surface liquid water (Porco et al 2006), decomposition
of clathrates at depth (Kieffer et al 2006), and sublimation at depth due to frictional heating (Nimmo et al 2007) have all been proposed as plume generation
mechanisms.
The surface fractures radiate ∼15 GW (Howett et al 2011) and the plume
latent heat carries away ∼1 GW, and this energy must be continually resupplied
by some heat source. Characterization of the plume is crucial to understanding
the mechanism that maintains their activity. Currently ice particle production and
escape rates are not well constrained because of limited knowledge of the plume
particle size distribution (Porco et al 2006; Spahn et al 2006b). However, the mass
production rate of the plume gas has been estimated to be about 120-180 kg/s
from occultation data (Tian et al 2007). This value is surprisingly high, sufficient
to remove a significant fraction (∼20%) of Enceladus’ mass over the age of the
solar system (Kargel 2006). These tiger stripes are typically ∼500m deep, ∼2 km
wide, up to ∼130 km long and spaced ∼35 km apart. Crosscutting fractures and
ridges with almost no superimposed impact craters characterize the area between
the sulci. The tiger stripes are apparently the source of the geysers observed to
emanate from the south polar region (Spitale and Porco 2007; Spencer et al 2009).
Tectonic features dominate the surface, demonstrating, that also smooth terrain is
pervasively tectonized (see section 2.1.4). Several different tectonic processes seem
to have been at work. The sinuous chain of scarps that bound the south polar
terrain at a latitude of ∼55◦ appear to have formed in response to compressional
forces, while north-trending fracture zones that radiate from peculiar Y-shaped
cusps and interrupt the chain of scarps appear to be extensional (Porco et al
2006). Shear offsets along pre-existing rifts are also observed near the transition
between these contractional and extensional features.
While the south polar terrain is a focus of pervasive active tectonism, other
regions are not (e.g., the cratered north polar region). Analysis of the relationship between impact craters and tectonic features (Barnash et al 2006; Bray et al
2007) indicates that the tectonism has persisted through time. Furthermore, fossil terrains elsewhere on Enceladus reminiscent of the south polar terrain suggest
multiple resurfacing episodes throughout the satellite’s history (see section 2.1.4).
The chain of scarps bounding the south polar terrain, as well as the northwardradiating fracture zones, may be the product of a change in Enceladus’s global
figure. The coherent (but latitudinally asymmetrical) global pattern of deformation on Enceladus is currently unexplained, but it is most probably due to shape
Enceladus as an active body
17
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changes, perhaps related to a hypothetical episode of true polar wander (Nimmo
and Pappalardo 2006).
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2.3.2 Internal dynamics and reorientation
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Several observational constraints suggest that the ice shell of Enceladus is dynamically active, at least in the southern hemisphere. The existence of the young
terrain (tiger stripes) at the South Pole may be evidence for mobile lid tectonics,
perhaps associated with convection (Barr 2008). Similarly, the high polar heat
flux may be due to episodic overturn of a convecting ice shell (O’Neill and Nimmo
2010). Moreover, the localization of heat production by tidal friction at the south
pole may be attributed to a regional thermal anomaly above a liquid water reservoir (Tobie et al 2008). Outside the south polar terrain, several theoretical studies
showed that thermal convection is possible (Barr and McKinnon 2007; Roberts
and Nimmo 2008; Běhounková et al 2010). However, no observation permits us to
confirm the occurrence of convection on a global scale.
Stereo topography shows that the surface of Enceladus is notably bumpy at the
∼100 km lengthscale (Schenk and McKinnon 2009) (see section 2.1.1). The ∼1 km
amplitude of some of the depressions is too large to be plausibly due to thermal
convection, even in the event of ice shell yielding (Showman and Han 2005). On
the other hand, shell density or thickness variations could easily explain such
amplitudes. Shell thickness variations would be rapidly erased if convection takes
place within the shell, but could persist for longer periods in a conductive shell
(Collins and Goodman 2007). Density variations in the cold, near-surface lid could
potentially co-exist with deeper convection, but would need to be large (δρ/ρ ∼ 0.1)
to explain the topography anomalies. Perhaps the simplest explanation for the
observed topography is that the ice shell of Enceladus is static and conductive,
except near the south pole.
The polar location of the hot spot is plausibly due to reorientation of the
satellite driven either by a low-density blob (diapir) within the silicate core or the
ice shell (Nimmo and Pappalardo 2006), or by the topographic low at the south
pole (Collins and Goodman 2007). In the diapir case, motion is not necessary;
all that is required is a density contrast. Ice shell density contrasts needed for
reorientation are large, more suggestive of compositional variations than thermal
variations. However, compositional variations could be induced by melting caused
by thermal variations (Pappalardo and Barr 2004; Tobie et al 2008). Studies of
coupled thermal-compositional convection on Enceladus are currently in their infancy (Stegman et al 2009). It has been suggested that the global tectonic pattern
on Enceladus might be the result of a ∼90◦ reorientation episode (Matsuyama and
Nimmo 2008). However, the spatial density distribution of crater does not seem
consistent with recent polar wander (Kirchoff and Schenk 2009). More detailed
geological mapping, summarized in Spencer et al (2009), lends some credence to
this idea, although another mechanism (such as diurnal tides) has to be invoked
to explain the apparently tensional nature of the tiger stripes. The best evidence
that reorientation is likely to have occurred would be the detection of a negative
gravity anomaly centred on the south pole (Nimmo and Pappalardo 2006).
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2.3.3 Ocean lifetimes
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The lifetime of any subsurface water depends on how rapidly heat is being extracted into the overlying shell, and how much heat is being added to the water
either from internal dissipation or by the silicate core beneath. Because there is
expected to be little heating in the core, and the shell of Enceladus is relatively
thin, predicted global ocean lifetimes are only a few tens of Myr (Roberts and
Nimmo 2008).
A simple analytical treatment illustrates the problem. In steady state, the
temperature at the base of a spherically-symmetric, bottom-heated conductive
shell is given by:
Tb − Ts =
Fb
(R − f R )
k
(1)
18
Sascha Kempf et al.
664
where Ts is the surface temperature (80 K), Fb is the heat flux at the base of the
shell, k is the thermal conductivity (assumed constant), and R and f R are the
outer and inner shell radii, respectively.
If tidal heating is concentrated in the warmest, basal ice (or in the ocean below;
Tyler (2011)) then equation (1) will provide a reasonable estimate of the basal ice
temperature (assuming spherical symmetry). Heating distributed throughout the
shell would result in a lower temperature, and convection will only increase the
rate of heat loss, so this approach is conservative. The steady state tidal heating
contribution is about 1.1 GW ( assuming Q = 18 000 in Saturn, Meyer and Wisdom
(2007)), resulting in a basal heat flux of about 4 mW m−2 for f = 0.6; to this
must be added the silicate contribution of about 1 mW m−2 (Schubert et al 2007).
Application of equation (1) and taking k=3 Wm−1 K−1 then results in a basal
temperature of about 180 K. This result shows that steady-state tidal heating
irrespective of where it is being generated – is too small to permit a global ocean
to survive, except in cases of extreme melting point depression (the ammonia-water
eutectic freezes at 177 K).
On the other hand, a regional sea can be sustained. Although equation (1) will
only be approximately correct if the heating varies spatially, it does suggest that
increasing the local basal heat flux by a factor ∼2 would be sufficient to maintain
a local water body in steady state. Spatial variations in tidal heating are likely at
Enceladus (e.g. a warm blob at the south pole will focus dissipation), and detailed
numerical modelling indicates that a regional sea can indeed persist in quasi-steady
state (Tobie et al 2008; Běhounková et al 2010).
In summary, a long-lived regional body of water seems more likely than a shortlived global ocean. If complete freezing ever takes place, dissipation in Enceladus
as a whole will decrease significantly, and equation (1) shows that it is then very
difficult to re-melt the ice.
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2.3.4 Coupled thermal-orbital evolution and episodic activities
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A critical constraint on the evolution of Enceladus is that the observed current
rate of heat loss significantly exceeds the rate (1.1 GW, assuming Q = 18, 000
in Saturn) which can be maintained in steady state by tidal dissipation. Other
present-day sources of heat (e.g. radioactivity) are small, and primordial sources
(e.g. accretion, 26 Al decay) will have been lost long ago.
There are thus only two possibilities: Enceladus is generating heat at the
steady-state rate, but the rate of heat loss varies with time; or heat production
itself varies with time. Either eventuality means that we are seeing Enceladus at
a special time.
Episodic heat loss can arise as a natural consequence of convection, associated
to episodic resurfacing events (Tobie et al 2008; O’Neill and Nimmo 2010). Fig 5a
shows the surface heat flux for a convecting fluid with a constant internal heating
rate and a relatively weak lid, as predicted by O’Neill and Nimmo (2010). Every so
often, the convective stresses become large enough to cause localized yielding and
downwelling of the lid, bringing warm ice to the surface and resulting in a pulse
of heat. The heat pulses last ∼10 Myr and are separated by intervals of quiescence
of several hundred Myr. In this case, the heat production rate is the steady-state
value and thus Enceladus’ eccentricity need not have changed at all. This scenario,
although undoubtedly not correct in detail, thus provides one way of explaining
Enceladus’ evolution.
Alternatively, Enceladus may have undergone oscillations due to the coupling
between eccentricity, tidal heating and thermal evolution (cf. Hussmann and Spohn
2004). However, the eccentricity damping timescale is fast compared to the likely
timescale for heat transfer across Enceladus. Thus, the kind of oscillatory behaviour proposed by Ojakangas and Stevenson (1986) for Io does not occur at
Enceladus (Meyer and Wisdom 2008), irrespective of whether the heat transfer is
convective or conductive. Other possibilities do exist, however. For instance, Fig 5b
shows a scenario similar to that proposed by Stevenson (2008). Here Enceladus is
initially in a non-dissipative state, so the eccentricity increases; once the resulting
Enceladus as an active body
19
Fig. 5 a) Heat flux as a function of time for a 2D Cartesian convection model undergoing
episodic localized overturn events (O’Neill and Nimmo 2010). b) Heat production and eccentricity variations with time for a case in which shear heating is initiated at one eccentricity
value and ceases at a second value (see text).
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diurnal tidal stresses exceed some critical value, fault motion and shear heating
initiate, and dissipation then drives the eccentricity back down. Once a different
threshold value is reached, fault motion ceases and the cycle repeats. In this case,
both the eccentricity and the heat production vary with time. Similar to Fig 5a,
the heat pulses are short compared to the intervals of quiescence.
How may we distinguish between these two possibilities? One obvious difference
is that Fig 5b requires the eccentricity to vary with time. The rate of eccentricity
damping at Enceladus is given by (e.g. Sohl et al 1995):
−a(1 − e2 )
de
=
P
dt
GMS MT tide
(2)
723
where e is eccentricity, a is semi-major axis, Ptide is the tidal dissipation rate and
MS and MT are the masses of Saturn and Enceladus. For a present-day dissipation rate of 15 GW, the eccentricity damping timescale is about 0.8 Myr and the
timescale for the change in mean motion is about 7.5 × 10−11 yr−1 . Such rates are
potentially detectable, given a sufficiently long baseline of observations as it was
done for Io (Lainey et al 2009). However, the change in eccentricity and mean motion depends also on the dissipation inside Saturn, which counterbalance the effect
of Enceladus’ dissipation. Preliminary analyses by Lainey et al (2010) indicates
that tidal dissipation in Saturn is about ten times larger that what was initially
anticipated. If this is confirmed, it would indicate that Enceladus may dissipate up
to 10-12 GW without damping its eccentricity. This would imply that Enceladus is
close to equilibrium tidal heating, and that episodic activity is no longer needed.
Clearly, understanding the long-term evolution of Enceladus is a difficult problem, involving a non-spherically-symmetric body in which the thermal and orbital
behaviour are tightly coupled. So far relatively little work has been done on this
topic. One exception is Zhang and Nimmo (2009), which examined the orbital evolution of Enceladus and Dione and concluded that a convective Enceladus would
be too dissipative to permit the present-day characteristics of the Enceladus:Dione
resonance. However, all the previous works on orbital evolution would need to be
revised if high dissipation inside Saturn is confirmed. This is an area in which
future work is likely to prove fruitful.
724
3 Geochemstry
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3.1 Chemical observations
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3.1.1 Surface composition
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The surface of Enceladus is mainly composed of water ice. This is evidenced from
observations at the visible and near infrared (0.8-5.1 µm, Cruikshank (1980); Bu-
20
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763
ratti and Veverka (1984); Cruikshank et al (2005); Emery et al (2005); Brown et al
(2006); Verbiscer et al (2006); Newman et al (2008); Filacchione et al (2010), and
far ultraviolet (115-190 nm, Hendrix et al (2010)) wavelengths. The exceptionally
high geometric albedo of the moon (Buratti and Veverka 1984; Verbiscer et al
2005) is consistent with these interpretations. The dominance of water ice grains
in the plume’s solid emissions (Porco et al 2006; Hedman et al 2009; Teolis et al
2010; Postberg et al 2011) and the E ring (Hillier et al 2007; Postberg et al 2008,
2009) implies that the moon’s surface is partially covered by a thin layer of icy
grains emitted from the subsurface. The deposition of icy grains on Enceladus
surface is observed and described (Kempf et al 2010; Schenk et al 2011; Degruyter
and Manga 2011). Disc-integrated spectra of the moon indicate the dominance of
crystalline water ice (Grundy et al 1999; Emery et al 2005).
Photometric and spectral analysis of Cassini near infrared reflectance spectra
also reveals the dominance of crystalline ice except the areas between linear tectonic depressions (the tiger stripes) at the south polar region (Brown et al 2006;
Newman et al 2008). These local areas demonstrate the largest signature of amorphous ice. However, a high degree of ice crystallininity is observed at the warm
tiger stripes and could be caused by crystallization of amorphous ice (Newman
et al 2008).
Other surface compounds are not abundant and mainly imbedded in water
ice. Traces of CO2 complexed with ice and light organic compounds at the tiger
stripes are reported based on Cassini near infrared observations (Brown et al 2006).
Free CO2 is also seen northward of the tiger stripes region (Brown et al 2006).
Ammonia and/or ammonia hydrates have been suggested as possible trace (< 1%)
impurities based on interpretation of near infrared (Emery et al 2005; Verbiscer
et al 2006) and ultraviolet spectra (Hendrix et al 2010). These interpretations are
consistent with the detection of CO2 , NH3 , and organic species in plume gases
(Waite et al 2006, 2009, 2011). The occurrence of high molecular weight organic
molecules in surface materials agrees with the likely presence of those species
in the plume and E ring particles (Postberg et al 2008). Correspondingly, the
dominance of salt-bearing icy particles in solid plume emissions (Postberg et al
(2011); Section 3.1.3) implies the presence of salts in the south polar surface layer
composed of deposited plume particles. Re-accretion of the plume and E ring saltbearing particles (Section 3.1.3) also suggests a global occurrence of Na-, K-, Cl-,
and carbonate species in the surface materials.
764
3.1.2 Composition of plume gases
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The composition of the plume gases has been measured in situ by the Ion and
Neutral Mass Spectrometer (INMS) (Waite et al 2004, 2006, 2009) and remotely
by the Ultraviolet Imaging System (UVIS) (Esposito et al 2004; Hansen et al 2006,
2008, 2011). In addition, particle measurements by the Cassini Dust Analyzer
(CDA) in the E ring and in the plume (Srama et al 2004; Postberg et al 2009,
2011), the near-Enceladus plasma environment (Smith et al 2005, 2007, 2008;
Christon et al 2013) and surface composition (Brown et al 2006; Hodyss et al
2009; Hendrix et al 2010) provide indirect evidence of the plume’s composition.
The first measurements of plume composition were made with the INMS and
UVIS instruments during the second targeted Enceladus flyby (E2 Tab. 1) on 14
July 2005 and revealed that the plume is composed predominantly of water. The
E2 flyby was an upstream pass, during which the Orbiter only grazed the plume’s
low-density outer edge, with a minimum distance of 330 km to the plume sources.
INMS pointing during the pass was not favorable to achieving the maximum signalto-noise ratio. Despite these less than ideal conditions, however, INMS was able to
identify, in addition to H2 O as the dominant species, significant amounts of CO2
and CH4 in the plume as well as an unidentified species with a mass-to-charge
(m/z) ratio of 28 (Waite et al 2006), which was initially thought to be either CO
or N2 . The INMS E2 data also indicated the possible presence of trace quantities of
C2 H2 , C3 H8 , NH3 , and HCN. Mixing ratios relative to H2 O for the most securely
identified species are: CO2 0.6% (+/- 0.15%), CH4 0.23% (+/- 0.06%), and
Enceladus as an active body
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NH3 0.7% (+/- 0.2%). These values are determined from INMS measurements
made during the low-speed, low altitude Enceladus flybys (E7, E14, E17, E18,
cf. Tab. 1), which are thought to be the measurements most representative of the
plume composition (Magee et al., paper in preparation, 2013).
UVIS occultations of Enceladus’ plume have yielded compositional results
complementary to the INMS measurements. The first UVIS measurement of the
plume came from the stellar occultation occurring on July 14, 2005 in conjunction
with the E2 encounter. Absorbance detected in the far ultraviolet (FUV, 111.5191.4 nm) channel confirmed H2 O dominance in the plume gas (Hansen et al 2006).
The FUV channel offered a good opportunity to observe the presence of CO and
help clarify the INMS ambiguity at 28 Da. The measured absorption spectrum,
however, showed no sign of CO’s known absorption bands, leading Hansen et al
(2006) to set an upper limit mixing ratio of 1%. The second UVIS plume observation was obtained during a stellar occultation on October 24, 2007 where
absorption spectra in the FUV channel again failed to show signs of CO (Hansen
et al 2008). The non-detection of CO and formal 2σ upper limit mixing ratio of
3% relative to H2 O set by Hansen et al (2008). CO was also not detected in the
Visible and Infrared Mapping Spectrometer (VIMS) surface reflectance spectra
from Enceladus’ south polar region (Brown et al 2006). The most recent UVIS
observation of Enceladus plume comes from the solar occultation event on May
18, 2010. The extreme ultraviolet (EUV, 55-110 nm) channel was utilized for this
observation, allowing for possible detection of N2 . The UVIS absorption spectrum
showed that that N2 absorption feature was limited to a level corresponding to
the signal standard error, constraining the abundance of N2 to an upper limit of
0.5% (Hansen et al 2011). Absorption cross-sections are temperature sensitive, so
quantitative abundance determinations of other species from UVIS data would
require a full spectrum absorption cross section for H2 O below 200 K.
Analysis of high-sensitivity INMS spectra from the high-speed E5 flyby, during
which the Orbiter traveled closer to the plume center on a north-south trajectory
following the flow direction of the plume (Tab. 1), suggested that the m/z 28 signal
was due to CO produced in the detector by the hypervelocity impact dissociation
of CO2 plus a smaller amount of either N2 or C2 H4 ; any contribution to the signal
from endogenous CO was considered to be small and indistinguishable from the
dissociation product (Waite et al 2009). However, analysis of data acquired during
the low-altitude, low-speed E7 pass directly through the plume center (Tab. 1)
suggested that CO, if present, cannot be fully explained by CO2 impact dissociation and any CO native to the plume (Waite et al 2011). According to one of the
compositional scenarios discussed by Waite et al (2009), if N2 is assumed to be
present in the plume, then HCN must also be present. However, the detection of
HCN reported by Waite et al (2011) based on the E7 data has not been confirmed
by data from subsequent low-altitude flybys; within statistical uncertainties, the
m/z 28 signal can be accounted for by several combinations of species (e.g., CO,
N2 , HCN, C2 H4 ) (Waite et al (2009), Supplementary Information, and Magee et
al., paper in preparation, 2013).
N + detected by the Cassini Plasma Spectrometer (CAPS) in Saturn’s inner
magnetosphere (Smith et al 2005) could in principle indicate the presence of N2
(as well as of NH3 ) in the plume (Smith et al 2007). There has been no definitive
detection of the N2+ parent ion, however, and the CAPS data permit only an upper
limit to be established for it (Smith et al 2008).
Suprathermal mass-28 ions have been detected with the CHarge-Energy-Mass
+
Spectrometer (CHEMS), but the contributing species—whether N+
2 , CO , or some
other species—-cannot be resolved (Christon et al 2013).
Based on the data from the E5 flyby, Waite et al (2009) reported the definitive
detection of NH3 and the measurement of a deuterium-to-hydrogen ratio in the
plume H2 O close to the value determined for a number of comets (but see also
Hartogh et al (2011)) and twice the terrestrial D/H ratio. The presence of C4 and
heavier hydrocarbons and the tentative detection of 40 Ar as well as of formaldehyde and methanol were also reported. However, INMS data from the E7 and later
flybys during which the Orbiter passed at low altitudes directly through the center
22
Sascha Kempf et al.
882
of the plume provided no evidence for 40 Ar and, moreover, showed substantially
reduced abundances for the majority of the trace species reported by Waite et al
(2009) (Figure 6). Spectral signatures for C3 hydrocarbons are clearly evident in
the low-altitude data sets, although there are ambiguities in the identification of
the individual C3 hydrocarbons because different combinations of species can be
assigned to a particular m/z value according to the compositional model selected
to fit the data. Only upper limits can be given for heavier hydrocarbons (such
as C6 H6 ). The E7 and later low-altitude flybys occurred at velocities relative to
Enceladus that were a little less than half that of the E5 flyby (Tab. 1). Waite
et al (2011) thus interpreted the compositional differences between the E5 and the
low-speed data sets in terms of a velocity effect involving the impact dissociation
of certain molecules at the higher kinetic energies of the E5 flyby. For example,
the signatures of C4 and heavier hydrocarbons seen in the E5 data but not in
the E7 data may indicate the presence of fragments of larger, more complex organic molecules (with masses outside the INMS 100-Da range) that break up upon
impact with the surface of the instrument’s antechamber.
Other lines of circumstantial evidence also point to the likely presence of organic macromolecules and simpler hydrocarbons in the plume. Mass-spectral analysis of E-ring particles believed to originate in Enceladus’ plume (Type II particles)
indicates that they contain silicates and/or fragments of refractory organics such
as “polar organic compounds (e.g., alcohols, nitriles, acids) or long-chain hydrocarbons” (Postberg et al (2008); see also section 3.1.3). Additional evidence comes
from surface reflectance data at various wavelengths (cf. Section 3.1.1). Absorption
features in the VIMS spectra at 3.44 µm and 3.53 µm have been interpreted as
signatures of short-chain organics present in the tiger stripes (Brown et al 2006).
Hodyss et al (2009) attributed the feature at 3.53 µm to methanol, which is either
produced photochemically on the surface of Enceladus or released in the plume
and deposited on the surface (INMS data from the slow flybys allow only a conservative upper limit of < 0.004% to be established for methanol [Magee et al., paper
in preparation, 2013].). Hendrix et al (2010) found that the best fit to spectra of
Enceladus’ surface at FUV, middle-ultraviolet, and visible wavelengths is achieved
with spectral models that combine small amounts of NH3 and tholin with H2 O. An
H2 O-NH3 -tholin model spectrum was also recently used by Zastrow et al (2012)
to fit a composite Enceladus reflectance spectrum generated from Cassini UVIS
and HST COS and STIS data. To account for the possible presence of tholins
in Enceladus’ ice, Hendrix et al (2010) suggested that the “complex hydrocarbons” reported by Waite et al (2009) “could undergo processing to result in the
tholin-type material predicted to be present.”
883
3.1.3 E ring and plume particles
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The Saturnian system is swamped by small particles (size < 10 µm) escaping from
Enceladus gravitational domain, most of them confined within Saturn’s outermost
diffuse E ring. The Cosmic Dust Analyser (CDA) onboard the Cassini spacecraft
allows for chemical characterization of these particles. The instrument produces
time-of-flight mass spectra from cations and cationic aggregates of the gas and
plasma cloud generated by high-velocity impacts of single grains onto a metal
target (Srama et al 2004). CDA carried out compositional measurements of both
E ring particles (Hillier et al 2007; Postberg et al 2008, 2009) and freshly ejected
particles during plume crossings (Postberg et al 2011). Another dust family of
Enceladus origin are the Saturnian stream particles (Kempf et al 2005; Kempf
et al 2005; Hsu et al 2010, 2011a,b). These are tiny nanometer-sized grains whose
dynamics are not governed by gravity but by electromagnetic interaction. Spectra
recorded in the E ring are best for compositional analysis, as the spacecraft spent
much more time recording spectra while crossing the E ring than in the plume.
Thus, E ring spectra have an excellent statistical basis. Moreover, the dust density
during plume crossings resulted in such extreme impact rates that the stress on the
instrument led to disturbances of the spectral signal, lowering the compositional
information of individual grains.
Enceladus as an active body
23
Fig. 6 The accumulated INMS signal mass spectra for E5 and E7 encounters of the Enceladus plume representative of the high and low velocity distributions. The spectra have been
normalized so that the signal in other masses can be compared to the H2 O signal at 18 Da.
The lower bound of the plot window represents the noise level for both data sets. A significant
dependence on spacecraft velocity is observed when comparing the two spectra. The E5 spectrum shows composition with complexity up to C6 H6 (78 Da) whereas E7 complexity peaks
in the C3 group (around 40 Da). (Note: These spectra are presented merely to illustrate the
velocity effect observed in the INMS data and should not be used for detailed analysis. New
spectra and mixing ratios based on an extensive re-analysis of the INMS plume data from all
of the relevant Enceladus flybys are currently in preparation.)
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E ring grains Even before Cassini’s arrival it was suspected that water plays an
important role in the E ring’s composition. However, only Cassini revealed that
Saturn’s environment outside the main rings is totally governed by water ice particles (Hillier et al 2007) as well as water molecules and ions which form from water
dissociation (Young et al 2005; Esposito et al 2005; Melin et al 2009). Almost
all particles encountered in the E ring are dominated by water ice but nevertheless there are significant differences in their individual composition. E-ring impact
spectra obtained with CDA can be categorized into at least three distinct families
(Postberg et al 2008, 2009). Type I spectra imply almost pure water ice grains
whereas Type II spectra exhibit contributions from organic and/or siliceous material. Type III spectra show a drastically increased alkali salt content. Although
there are a few cases where individual spectra show characteristics of two types,
they form distinct families.
Type I particles About 65% of all particles above the detection threshold belong
to Type I. The abundance of this type increases further for small particles sizes.
Mass lines caused by water cations (H2 O)n (H3 O)+ , n = 0 . . . 15 are dominant or
exclusive. Na+ and K+ and their respective water cluster ions (H2 O)n (Na, K)+
are present in most cases and form the only non-water mass lines. Calibration
experiments point at very low concentrations of alkali salts with an Na/H2 O
ratio in the order of 10−7 (Postberg et al 2009).
Type II particles Type II spectra represent the second most abundant E ring
family and show the same characteristic as Type I with an additional distinct
feature between mass 27 and 31 which in most cases represents more than one
mass line (Postberg et al 2008). In many cases additional non-water signatures
appear and the peaks are broader than in the case of Type I or III. The
additional peaks strongly point at unspecified organic compounds with strongly
varying concentration. A contribution of silicates is also possible. However,
water ice is always the dominant species.
Type III particles This family of about 6% of E ring spectra exhibit a totally
different pattern of mass lines. In contrast to Type I and II, the water cluster
peaks (H2 O)n (H3 O)+ are absent or barely recognisable. The characterizing
mass lines are of the form (NaOH)n (Na)+ , n = 0 . . . 4 indicating a Na/H2 O
24
Sascha Kempf et al.
Fig. 7 Composition and size profile of the Enceladus’ ice particle plume (modified after Postberg et al (2011)). The background colours on the left panel show the fraction of salt-rich
grains. Overlaid are contours of constant mean particle radius obtained from the model. The
projection used is in the plane of the spacecraft trajectory during Cassini’s E5 flyby (solid
black line, shown with 10 s intervals). CDA measurements show that both the largest particles
and the highest fraction of salt-rich grains a few seconds after closest approach to Enceladus.
Structures of the three most relevant localized supersonic jets for this projection are clearly
visible in both the compositional profile and size contours. Note that the model only considers
particle sizes above the instrument’s detection threshold (0.2 m) and it shows the plume emission with an added E ring background that is increasingly relevant north of the south polar
source region. The right panel shows the modeled contribution from the few collimated supersonic jets versus more numerous slow, diffuse sources distributed all along the tiger stripes.
Most importantly it shows the fraction of salt-rich ice as a function of logarithmic altitude
above the south pole. Near the surface salt-rich particles account for about 70% of all grains
> 0.2µm but due to their larger average size they account for almost all the solid mass created
by Enceladus’ active region.
942
mole ratio above 10−3 . Frequent mass lines of NaCl − Na+ and Na2 CO3 − Na+
reveal NaCl and NaHCO3 + and/or Na2 CO3 + as the main sodium bearing
compounds. Calibration experiments indicate an average concentration of 0.5 −
2% sodium and potassium salts by mass with K compounds being less abundant
by a factor of at least 100 (Postberg et al 2009). The composition has a very
good match to the composition predicted for liquid water on early Enceladus
(Zolotov 2007) and implies a basic pH of 8 − 9. Impacts of Type III particles
have an average ion yield which is several times higher than of Type I, implying
a considerably larger size.
943
Plume particles In 2008 and 2009 CDA carried out compositional analyses on three
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plume crossings, E4, E5, and E7. The majority of spectra can be assigned to
one of the three types previously detected in the E ring. Whilst the number of
obtained plume spectra was limited to fewer than twenty on E4 and E7, a specific
configuration of CDA at E5 provided a spatial resolution high enough to infer a
compositional profile of the plume (Fig. 7).
The amount of salt rich Type III grains increases above 40% at the closest
approach to Enceladus’ surface (21 km). The observed compositional profile can
be reproduced when Type III grains are ejected with slower speeds corresponding
to their larger size. The best match points at about twice the average size of Type I
grains (Postberg et al 2011). Applying a size dependent stratification of the plume
(Schmidt et al 2008), which is in agreement with Cassini observations (Kempf
et al 2008; Hedman et al 2009), about 70% of all ejected particles are of Type III.
Owing to their larger size about 99% of the emitted solid mass must be salt rich
ice (Fig. 7). Naturally the small and fast salt-poor particles preferentially escape
the gravitational domain of Enceladus and therefore dominate the ice-population
of the E ring.
Enceladus as an active body
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25
The fit could be further improved by assuming sources with drastically varying
gas speeds. Hundreds of slower sources located all along the tiger stripes with gas
speeds of about 400 m s−1 would preferentially produce salt rich grains. In contrast,
the few collimated supersonic jets (Spitale and Porco 2007) observed in the UV
and the optical wavelengths with gas speeds above 1000 m s−1 (Hansen et al 2008,
2011) are enriched with the smaller, sodium-poor particles (Postberg et al 2011).
Furthermore, the CDA plume measurements imply a general enrichment of Type
II ice grains containing organic material in the plume, probably associated with
the fast collimated jets. Grains which are both enriched in salts and organic species
are frequent.
Saturnian stream particles Cassini encountered these tiny (2 − 8 nm in radius),
high-speed (> 100 km s−1 ) particles in 2004 even before its orbit insertion at Sat-
984
urn (Kempf et al 2005). The initial analysis of CDA impact spectra indicated
little, if any, water ice and a major siliceous component (Kempf et al 2005). More
surprising is the recent discovery that most stream particles start their dynamical
life in the water dominated E ring (Hsu et al 2011a,b). This points at a stream
particle release as refractory remnants of sputtered icy E ring grains in which they
were initially embedded (Hsu et al 2011b). A thorough revision of stream particle
spectra identified silica (SiO2 ) as the main component of most stream particles.
The most plausible explanation for this finding is the abundant formation of nanocolloidal silica in an aqueous phase which is then partly integrated in ice grains
expelled by Enceladus’ plume (Hsu et al., 2013, in preparation). Small amounts of
silica still embedded in ice grains with a Si/H2 O ratio below 10−4 would probably
go undetected by CDA. However, occasional mass lines at 28 Dalton in spectra of
E-ring and plume particles could be an indication for embedded nano-silica grains.
985
3.2 Composition of the interior
986
3.2.1 Bulk composition and a rocky core
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The density of Enceladus of 1608 ± 5 kg m−3 (Porco et al 2006) implies the presence of a rocky component in the interior. For the assumed densities of ice with
some contaminants (1010 kg m−3 ) and the rock (2500-3528 kg m−3 ), the rock mass
fraction is 0.62 to 0.52, respectively (Schubert et al 2007). In these evaluations, a
rock density of 2500 kg m−3 roughly represents either the grain density of highly
aqueously altered CI carbonaceous chondrites (Consolmagno et al 2008) or hydrated silicates, and the value of 3528 kg m−3 corresponds to the density of Io. A
low rock density is likely because aqueous alteration is supported by the detection
of salt-bearing grains emitted from the moon (section3.1.3).
The rocky (inorganic) part of the moon could have the composition of CI carbonaceous chondrites very similar to that of the solar photosphere. However, the
detection of typical cometary volatiles (CO2 , NH3 , and CH4 ) in the plume, the
comet-like abundances of NH3 and CH4 , and D/H ratio in plume gases imply an
enrichment of Enceladus in volatile and organic species compared to chondritic
materials. Comets are strongly enriched in light elements relative to CI carbonaceous chondrites and the C-H-O-N type organic materials account for ∼23 wt%
of the comet Halley dust (Jessberger et al 1988; Hanner and Bradley 2004). The
likely presence of high molecular weight organic compounds in the plume and Ering particles (sections 3.1.3 and 3.1.2) also agrees with an organic-rich interior of
the moon. Thus the composition of Enceladus’ non-icy solids could be closer to
the solar photosphere than to CI carbonaceous chondrites, which are depleted in
C, O, and N. The density of hydrated organic rich rocks with the solar abundance
of C could be ∼ 1900 ± 100 kg m−3 (Zolotov 2009).
In addition to abundant organic matter, the bulk density of rocks could be further lowered because of a high porosity. CI and CM carbonaceous chondrites have
a porosity of 35% and 23%, respectively (Consolmagno et al 2008), and the pressure in the moon’s interior (<∼ 0.5 kbar) may not be sufficient to compact these
26
Sascha Kempf et al.
Fig. 8 Thickness of the icy shell as a function of core density for a completely differentiated
Enceladus (modified after Schubert et al (2007)). The error bars correspond to core density
of 2500-4000 kg m−3 for hydrated and anhydrous rocks, respectively. An organic rich, porous,
and hydrated core could have density below 2500 kg m−3 . Differentiated Enceladus with such
a low-density core could have a thin icy shell, as illustrated by arrows.
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rocks below 10% porosity (Zolotov 2009). By analogy with chondrites (Consolmagno et al 2008) this porosity could be mostly accounted for by microporosity
so that the pores may not be filled with water and/or ices. If this is the case, it
follows that a hydrated organic-rich rock with the solar abundance of C and 10%
porosity could have a density of only ∼ 1700 kg m−3 . This low end-member density
corresponds to the rock mass fraction of 0.91. In a completely differentiated interior such a low density of the rocky core would correspond to a fairly thin icy shell
(Fig. 8). Depending on the core density, the ice shell may be as thin as ∼ 13 km.
The low upper limit for N2 in plume gases (Hansen et al 2011), a preservation of typical cometary species, and the low K/Na ratio in salt-bearing icy grains
(Postberg et al 2009) indicate predominantly low-temperature (T ≤273 K) waterrock interactions in Enceladus history (section 3.1). This implies a low amount of
accreted 26 Al and a formation of Enceladus at least several Ma after the Ca-Al-rich
inclusions (CAIs), in agreement with recent formation models for the Saturnian
system (Sasaki et al 2010). The low-temperature rock alteration is in agreement
with the global shape of Enceladus which is not consistent with hydrostatic equilibrium (Porco et al 2006; Thomas et al 2007). The shape suggests that the rocky
core has never reached temperature high enough to relax core topography. In this
condition, a large porosity may have persisted in the organic-rich and hydrated
core, as discussed above.
Although some minor silicates may be present in icy particles emitted from
Enceladus (Postberg et al 2008), the data obtained with CDA are not sufficient to
evaluate composition of Enceladus’ rocks. However, the supposed low-temperature
alteration of the rocks implies that CM2 and some other carbonaceous chondrites
(e.g., Zolensky and McSween (1988); Brearley and Jones (1998); Brearley (2006)
could be used as a proxy for an altered mineralogy of the core. Inorganic solids
could consist of abundant phyllosilicates (serpentine, cronstedtite, saponite, chlorite), primary and secondary sulfides (troilite, tochilinite, pyrrhotite, pentlandite),
magnetite, phosphates, and carbonates. Many of these minerals are modeled to be
present in an aqueously altered Enceladus’ core (Zolotov 2007). The organic fraction of the rock could consist of aqueously altered cometary-type C-H-O-N compounds (Hanner and Bradley 2004) and may also contain high molecular weight
polymers abundant in carbonaceous chondrites (Septhon 2002). Water-soluble and
light organic compounds as well as high-solubility Na-rich salts could concentrate
at the ice-core boundary along with NH3 and methanol that strongly decrease the
melting temperature of ice (sections 3.2.3 and 3.1). CO2 from melted ices converts
Enceladus as an active body
27
0.10
CO
0.1
H2 S
200*
1
2
CH
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N2
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Pressure (MPa)
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Unstable
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Fig. 9 Stability curves for clathrate hydrates detected or suspected at Enceladus. The dashed
curve corresponds to the mixed CO2 -CH4 clathrate composition at the CO2 /CH4 =4. The
red curves show modeled temperature profiles in the thermally conducting icy shell for different heat flow values (5 to 200 mW m−2 ). The curve labeled 200* corresponds to the surface
temperatures of 200 K. Other red curves are for surface temperatures of 73 K. The red domain
indicates the pressure-temperature conditions in a putative convecting ice layer. A temperature
of 240 K is the typical value expected in cold downwelling, while hot upwellings typically have
temperature of the order of 260-270 K. The blue area shows the stability field of liquid water.
The scheme in the lower part of the plot explains the conditions under which the clathrates are
stable. The clathrate stability curves have been obtained with the CSMHYD software (Sloan
and Koh 1998), except for the curve of CO, which is determined from the August equation
provided by Hersant et al (2004) for T <273.15 K and from experimental data for T >273.15 K
(Mohammadi et al 2005).
1056
to carbonate/bicarbonate species at alkaline conditions created by rock alteration.
Organic species could be presented by low-viscosity compounds (including liquids,
section III.2.3). Both organic and water-based liquids could decouple the core from
the icy shell and favor tidal heating. High-solubility salts could be mainly presented
by alkali chlorides, carbonates, and bicarbonates (Zolotov 2007) detected in the
solid emissions (Postberg et al (2009, 2011); section 3.1.3). Sulfates are not detected
and they may not have formed in reduced low-temperature aqueous environments.
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3.2.2 Icy shell
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As already discussed in sections 2.1.3 and 3.1.1, the surface of Enceladus is mostly
composed of water ice, both crystalline and amorphous (Brown et al 2006). The
plume is also dominated by H2 O, in the form of vapor and ice particles (Hansen
et al 2008; Waite et al 2009; Postberg et al 2011). However, the observation of
non-water gas molecules in the gas plume (Waite et al 2009, 2011, section 3.1.2) as
well as the detection of salt-bearing grains emanating from the south pole region
(Postberg et al 2009, 2011, section 3.1.3) indicate the presence of impurities in the
icy shell.
The gas molecules may be dissolved in liquid water (in a subsurface reservoir
and/or localized melt pockets, e.g., Porco et al (2006), and/or trapped in the form
of clathrate structure within the icy shell (e.g., Kieffer et al (2006)). Clathrate hydrates are crystalline structures made of hydrogen-bonded H2 O molecules, which
form cages that can trap gas molecules (e.g. Sloan and Koh (1998); Lunine and
Stevenson (1985); Choukroun et al (2013)). The composition of common clathrates
is G × n H2 O, were G is a gas molecule and n ≈ 5.75. Clathrates occur in Earth’s
permafrost, oceanic seafloor sediments and could store volatiles in various solar
system bodies (see Choukroun et al (2013) for a review). On Enceladus, they have
been proposed to play a dominant role in storing the gases, driving the plumes
(Kieffer et al 2006) and controlling the tectonics by their stiff brittle nature (Gioia
et al 2007a).
28
Sascha Kempf et al.
Fig. 10 Concentrations of volatile species in the Enceladus plume compared to the Henry’s
law constants, which characterize solubility of gas species in pure water. The lack of clear
correlation of plume concentrations with solubilities implies a significant contribution of plume
gases from dry sources.
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Fig. 9 illustrates the dissociation curves of several clathrate-forming compounds
compared to typical temperature profiles expected in the icy shell, corresponding
to different heat flow values. For all gas compounds, the dissociation temperature
increases as a function of pressure (depth). H2 S clathrate is always stable even in
the warmest temperature profile, corresponding to the conditions expected in the
vicinity of the tiger stripes ( ΦS = 200 mW m−2 , TS = 200 K, were ΦS and TS
are the heat flow and temperature at the surface). The high stability of H2 S in
the clathrate phase may explain its low abundance in the gas phase of the plumes.
The other considered gas compounds are unstable in the form of clathrate for this
warm case. For 200 > ΦS > 50 mW m−2 , mixed CO2 − CH4 clathrates (having a
CH4 /CO2 ratio similar to the plume composition determined during the E7 flyby)
become unstable at depth comprised between 3 and 10 km. For TS = 180-200 K,
mixed CO2 -CH4 clathrate is always unstable. This implies destabilization of the
CO2 -CH4 clathrate in the vicinity of tiger stripes in the south polar region. Note
that destabilization occurs even without decompression caused by fissuring in the
ice. For ΦS < 50 mW m−2 , which may represent areas away from the south pole
region, clathrates of pure CO2 or CO2 -CH4 mixture are always stable. Clathrates
of CO and N2 are stable only in very cold regions and are unexpected within the
warm ice shell in the south polar region.
Despite the thermodynamic stability of clathrates, the role of clathrate dissociation as a main driver of plume activity remains a hypothesis. INMS data
indicate that the major clathrate-forming gases (CO2 , CH4 ) could compose less
than ∼1 mole % of the H2 O dominated plume. If these gases are released through
decomposition of clathrates plume source regions may contain less than 6 vol. % of
clathrates. Such a low clathrate volume fraction suggests that clathrate crystals
are mixed with ice crystals and probably do not form a spatially separated pure
clathrate reservoir.
Gases which do not form stable clathrates in the icy shell (e.g. H2 ) could be
present in defects of ice crystals and in the spaces between the crystals. NH3 , and
CH3 OH detected in the plume are very soluble in liquid water (Fig. 10) and do not
form clathrates. These molecules, especially strong anti-freezing species NH3 and
CH3 OH, are likely to be concentrated in a potential liquid phase (section 3.2.3).
The degree of compositional heterogeneity of the icy shell is unknown. The diverse nature of chemical species emitted from Enceladus may indicate spatial separation of salt- and volatile-bearing regions of the shell. Pure water ice zones could
Enceladus as an active body
29
1129
also be separated from salt-bearing water ices and volatile-bearing regions, which
may represent newer melted primordial icy materials. The composition structure
below the south polar region could be different from other regions. The compositional heterogeneity should reflect mechanisms of water-rock differentiation and
subsequent thermal processes in the shell (sections 3.1.1 and 3.1.2). In one possible
scenario, enriched liquids may be injected into the icy shell owing to the pressurization of a subsurface oceanic reservoir, resulting transient volume contraction
of the liquid reservoir upon crystallization (Manga and Wang 2007). The dikes of
enriched liquids may crystallize when reaching cold regions within the icy shell,
leading to local enrichment in oceanic impurities (salts, NH3 , methanol, soluble
and condensed organic species). Some oceanic materials could be concentrated in
frozen lenses (cf., Matson et al (2012)). In another scenario, salts and organic compounds could be abundant in the upper horizons of the icy shell reflecting either
ancient cryovolcanic activity or rapid freezing of early oceanic water after sinking
of a primordial icy-rocky shell (section 3.1.1). In all these cases, the icy shell probably consists of a mixture of ice with local enrichment clathrate hydrate(s), organic
compounds and salts concentrated upon crystallization of aqueous solutions.
1130
3.2.3 Possible liquid phases
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The stability of aqueous phases in the interior mainly depends on temperature
and concentrations of species that decrease the melting temperature of ice. The
energy output of the south polar region (>15 GW, section II.1.2) is sufficient to
melt significant parts of the south polar ice. Although Cassini chemical data do
not provide unequivocal evidence for large liquid water reservoirs, the occurrence
of aqueous solutions in the interior is likely.
The observed surface temperature of ∼190 K in the most active linear ridges of
the south polar region (Spencer et al 2011) is high enough to stabilize aqueous solutions even in surface materials. In fact, this temperature is higher than the temperatures of the ammonia-methanol-water (∼153 K), methanol-water (∼160 K), and
ammonia-water (∼175 K) eutectics (Kargel 1992). At the temperature of ∼190 K,
solutions rich in ammonia could exist as brine pockets in the surface ice. A high
NH3 /CH3 OH ratio in plume gases implies a dominance of ammonia in brine pockets.
The subsurface temperatures at the Tiger Stripes must exceed the surface
values (190 K) for any physically plausible scenario of heat-flow below the south
polar terrain. Moreover, properties of the observed flow of vapor and ice particles
(Schmidt et al 2008) imply subsurface temperatures well above 240 K. These elevated temperatures would correspond to less concentrated and more abundant
fluids in depth. The corresponding solutions could stably exist within ice as brine
pockets rich in neutral species (ammonia, methanol, NaCl) and ions (Na+ , Cl− ,
K+ , HCO3 − , CO23− , etc.) that both decrease the freezing temperature of an aqueous solution. The composition of fluid pockets should also reflect compositional
heterogeneity in distribution of volatiles and salts of the icy shell. Depending on
the origin of ices and temperature (sections 3.2.4; 3.3.1), the fluid pockets may be
rich either in dissolved salts or in NH3 . If ammonia is not locally present, the near
eutectic brines at the temperature ∼238-243 K could be rich in Na+ and HCO3 −
while warmer fluids are NaCl dominated and less saline solutions (Fig. 4 in Zolotov
(2007)). If the composition of the emitted ice grains reflects the composition of
the liquids they have an alkaline pH (Postberg et al 2009), which depends on the
temperature (Zolotov 2007).
The bottom of the icy shell is a place where current liquids are even more
likely. The rock-ice interface could be enriched in species that decrease ice melting temperature. The enrichment could have occurred through downward freezing
of a primordial ocean after differentiation (section 3.3.1). These impurities would
contain species detected in plume gases and particles. In addition, organic compounds extracted from rocks could be abundant (section 3.3.1) and may separate
in an organic liquid phase. The presence of these low-viscosity species would facilitate tidal motions and heating at the bottom of the icy shell (Zolotov 2007).
30
Sascha Kempf et al.
1189
Tidal motions should favor ice melting and will decrease the salinity of aqueous
fluids. In fact, a decoupling layer at the core-ice interface beneath the south polar
region is required to generate enough tidal dissipation to account for the observed
thermal emission (Nimmo et al 2007; Tobie et al 2008). Either a water reservoir
or a low-viscosity organic-rich layer at the core-ice interface could serve as a lubricant that favors tidal flexing. In any case, the temperature at the ice interface
needs to be close to the melting point of water ice to generate significant viscous
dissipation in the overlying ice shell. A large liquid water reservoir (ocean) – if any,
would be alkaline NaCl rich solution with subordinate amounts of CO23− , HCO3 − ,
and K+ ions (Zolotov 2007; Postberg et al 2009). Lower volume of the reservoir
would correspond to higher salinity. Glein and Shock (2010) used mass balance
calculations to evaluate NaCl content of this putative reservoir as a function of its
geometry (volume).
Despite the possible existence of pore space in the rocky core (section 3.2.1),
liquid phases may not exist within the rocky core. The depletion of long-lived radionuclides in rocks (Schubert et al 2007), the lack of major tidal motions within
rocks, and an accumulation of inorganic and organic impurities above the rocky
core (section III.3.1) do not favor stability of aqueous fluids within the core. However, water-based and organic liquids could occupy pore spaces in the uppermost
rock layer that underlie a liquid reservoir at the bottom of the icy shell.
1190
3.2.4 Aqueous vs. dry sources of plume emissions
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The striking compositional similarity of the E-ring salty ice grains (Postberg et al
2009) and solutions formed in water-rock interaction models (Zolotov 2007) indicates aqueous processes in Enceladus’ history. The dominance of salty ice grains
in solid plume emissions (Postberg et al 2011) indicates a major contribution
from aqueously processed materials. The question if the salty ice grains originate
from ices or fluids requires a thorough consideration. This section discusses chemical species detected at Enceladus in the context of phase composition of plume
sources.
The detected salt-rich plume grains (Postberg et al 2011) cannot originate from
a salt-poor matter and must be emitted from salty ices or aqueous solutions. Some
condensation of water on salty ice grains during their subsurface ascend implies
higher concentrations of salts in the source region(s) than in the salt-rich grains
forming the plume. The observed compositional uniformity of those grains (0.5-2
wt. % salt, Postberg et al (2009)) requires a fairly homogenous composition of the
source material, and plume sources should provide this homogeneity.
An aqueous source of salty ice grains is consistent with their compositional
homogeneity. In one possible scenario, an aerosol spray forms through bursting of
gas bubbles at the surface of subsurface aqueous reservoirs below back pressured
vents (Postberg et al 2009). The droplets freeze, thereby conserving homogeneous
composition of the fluid, and are dragged upward by the gas flow. The salty grains
formed from an aerosol spray would form larger ice grains than salt-poor grains
which form by nucleation from water vapour (Schmidt et al 2008). Therefore, this
scenario is consistent with the observed larger size of salt-rich grains if compared
with the salt-poor population (Postberg et al 2011).
Evaporation of water from unconnected subsurface reservoirs would unevenly
increase the fluid salt contents. However, even relatively small isolated reservoirs
in the order of a few kilometer equivalent radius would need tens of thousands of
years to change salt concentrations in the order of a few percent (Postberg et al
2009). If these time scales are considered for the currently observed plume activity
from subsurface, salty lakes’ are only consistent with large or connected water
reservoir(s) below the south polar region.
It is harder to explain compositional homogeneity in terms of dry sources
of salty grains. Sublimation of salty ices would lead to salt lag deposits at jet’s
source regions and the sublimation could not account for the larger size of salt-rich
grains. The earlier proposed origin of the plume’s icy grains through condensation
after sublimation (e.g., Nimmo et al (2007); Kieffer et al (2009)) is not consistent
Enceladus as an active body
31
Fig. 11 Concentrations of volatile species in the Enceladus plume (red symbols, see section
3.1.2) compared to the abundances in comets (blue boxes, data from Mumma and Charnley
(2011); A’Hearn et al (2012)). Note that CO2 has not been reported in some comets. Plume
concentrations of NH3 and CH4 closely match cometary data.
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with the detection of salt-rich grains. Suggested downward freezing of an early
salty ocean (see section 3.3.1) would have concentrated salts at the bottom of the
forming icy shell and may not cause homogeneous trapping of salts in the main
body of the icy shell. Even with flash freezing of early oceanic water it seems very
difficult to form large source regions of compositionally homogeneous salty ice.
Therefore, a comet-like violent sublimation which directly entrains ice particles in
the flow may not produce compositionally homogeneous salty ice grains.
The proposed formation of icy grains through an explosive decomposition of
clathrate hydrates (Kieffer et al 2006; Gioia et al 2007a) is not consistent with the
small fraction of clathrate-forming plume gases (mainly CO2 and CH4 ) indicating
a limited (<∼6 vol.%) amount of corresponding clathrates in plume source regions.
Another problem of this model is that water vapour is created by sublimation of
a major part of these ice particles which are initially entrained in the flow of gases
released from clathrates. The formation of sodium free water vapour (Schneider
et al 2009) through sublimation of icy grains in clathrate-decomposition models
is not consistent with the dominance of salt rich ice grains. Either some sodium
would enter the gas phase through grain sublimation or – if that can be avoided
– almost pure salt grains would be left behind in the gas flow which is also not
observed.
In contrast to the composition of ice grains the composition of plume gases does
not provide a strong constraint on liquid water as the major plume source (despite
the stability of NH3 /CH3 OH-rich solutions in the south polar ices). If plume gases
exclusively originate from a liquid water reservoir, low-solubility species would
be more abundant than high-solubility compounds. However, the most reliable
data show no clear correlation with solubilities in water (Fig. 10 ). On the one
hand, highly soluble methanol looks strongly depleted compared to cometary abundances. On the other hand, highly soluble NH3 and shows no signs of depletion,
whereas low-solubility CO, N2 , CH4 , C2 H2 , C2 H6 , and other hydrocarbons are less
abundant or not detected in the plume. The lack of a clear correlation of INMS
data with solubilities (Fig. 11) is not consistent with degassing solely from a subsurface liquid water reservoir. A low (far from eutectic) NH3 /H2 O ratio in the plume
implies a contribution from dry sources, though degassing of NH3 -rich brine pockets together with surrounding ices is possible. The detection of NaHCO3 and/or
Na2 CO3 in icy grains (Postberg et al 2009) indicates the presence of HCO3 − and
CO23− ions in the solution, and some CO2 degassing could be driven by changes
in solution pressure, temperature, or pH (Matson et al 2012).
The lack of abundant rocky solids in plume emissions put additional constrains
on the phase composition of plume sources. The dominance of ice in plume solid
emissions is consistent with separation of rocks during differentiation and does not
indicate a presence of cometary-type rock-ice mixtures in the plume source region.
32
Sascha Kempf et al.
Fig. 12 The K/Na atomic ratio in salt-rich E ring grains (Postberg et al 2009) in comparison
with modeled aqueous fluids on early Enceladus (Zolotov 2007, 2012). The plot indicates
predominantly low-temperature aqueous processes in the interior.
1281
The deficiency of rocks remains to be reconciled with the presence of CH4 , NH3 ,
among plume gases. One possible explanation is the heterogeneous accretion of
Enceladus (Charnoz et al 2011, section 3.3.1) that led to preferential accumulation
of solid volatile species (ices, clathrates, hydrates) in the peripheral parts of the
moon. Current tidal warming of the south polar region could cause both melting
of volatile-poor salty ices and release of gases from never melted volatile-rich ices.
Although the composition, grain size and dynamics of the emitted ice particles
strongly point at a source of salt-bearing aqueous solution it is not unequivocally
proven. The composition and abundance of volatiles in the gas plume is not in
agreement to be solely emitted by evaporation of water. However, all currently
proposed dry scenarios encounter conflicts with one or more of the plume observations. A mixture of liquid and solid sources contributing to the plume is the
likely scenario. Whereas the liquid sources would be responsible for most of the ice
emission, ice sublimation and possible clathrate decomposition have a significant
contribution to the gas phase.
1282
3.3 Chemical evolution scenarios
1283
3.3.1 Accreted materials and chemical effects of water-rock differentiation
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Enceladus could have accreted at low temperatures and pressures from either a
slow-inflow accretion disc in the solar nebula (Canup and Ward 2002, 2009) or
from massive rings of Saturn formed significantly later after the formation of the
planet (e.g., (Porco et al 2007; Charnoz et al 2010, 2011; Canup 2010)). Enceladus
was probably accreted as a mixture of ices, condensed organic compounds and
volatiles, and rocky components. Water ice dominated over other ices, and other
volatiles could have accreted as hydrates (NH3 , HCl), ices (CO2 , NH3 , CH3 OH),
and clathrates (CO2 , H2 S, CH4 , and some other light hydrocarbons) (e.g., Hersant et al (2008)). Although the nature of accreted organic matter is unclear, some
material could have resembled high molecular weight organics common in carbonaceous chondrites and C-H-O-N particles detected in the comet Halley dust (Jessberger et al 1988; Fomenkova et al 1994). The rocky component was presented by
anhydrous and reduced minerals and amorphous silicates, by analogy with primitive materials in chondrites (e.g., Brearley and Jones (1998)) and cometary dust
(Jessberger 1999; Hanner and Bradley 2004; Zolensky et al 2006). This component
may not have abundant chondrules. The minerals could be dominated by enstatite
(MgSiO3 ), forsterite (Mg2 SiO4 ), Fe-Ni-metal, and troilite (FeS).
If short-lived 26 Al was abundant in the rocks at the time of accretion, melting
of ices could have caused water-rock differentiation within ∼ 106 years (Schubert
Enceladus as an active body
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33
et al 2007). Hypothetic scenarios that included enhanced heat sources on early
Enceladus (due to abundant 26 Al and large tidal dissipation) may have involved
hydrothermal processes, some dehydration of earlier hydrated phases, intense alteration and oxidation of organic species, oxidation of NH3 and N2 production
(Matson et al 2007; Glein et al 2008). These pathways also suggest a significant
thinning of the primordial ice-bearing shell (Schubert et al 2007). However, 26 Al
could have significantly contributed to the thermal budget only if the satellite
accreted within a few millions years after the formation of the CAIs (the first
major solids formed in the solar nebula). This early accretion is difficult to reconcile with the present knowledge of the Saturn system formation (e.g., Sasaki et al
(2010)) and with the possible formation of mid-size satellite from Saturn rings (e.g.,
Charnoz et al (2011)). Typical cometary abundances of NH3 and CH4 in plume
gases (Fig. 11), the low limit for N2 in plume gases (Hansen et al 2011), and the
low K/Na ratio in E-ring particles (Fig. 12) also indicate moderate heating without major hydrothermal activity in Enceladus’ history. Without abundant 26 Al,
the differentiation could have been driven by the decay of long-lived radionuclides
and occurred about 0.5–1 Gyr after accretion (Schubert et al 2007; Malamud and
Prialnik 2013). The release of heat in mineral hydration reactions (e.g., Grimm
and McSween (1989); Malamud and Prialnik (2013)) contributed to water-rock
differentiation. Tidal dissipation in the warmed interior may have also accelerated
the differentiation process (Schubert et al 2007).
Independent of the source of heat, temperature rose more rapidly in the central part of the body and led to the melting of ice, resulting in the gravitational
separation of rocky grains and melts, which formed a growing liquid layer above
a rising rocky core (Schubert et al 2007). Initial low-temperature fluids were rich
in species that strongly decreased the melting temperature of ice (NH3 , methanol,
Cl− ) and became enriched in metals along with warming of fluids and dissolution
of minerals. Release of volatiles from melting ices led to diverse chemical transformations that included dissolution, hydrolysis, and other chemical interactions.
As an example, dissolution of C oxides could have led to carbonate, bicarbonate,
and formate (HCOO− ) anions. Hydrolysis of ammonia produces the ammonium
ion (NH4 + ) which reacts with anions leading to precipitation of NH4 Cl and ammonium carbonates/bicarbonates (Kargel 1992; Marion et al 2012). These reactions should have diminished the concentration of NH4 + in solution. Hydrolysis
of formaldehyde causes its polymerization to methylene glycol, CH3 (OH)2 . Hydrogenation of carbon species in H2 -rich environments (see next paragraph) may
lead to formation of formate, methanol, CH4 , and other hydrocarbons. These and
other chemical transformations should have changed the abundances of accreted
volatile compounds and led to new species that accumulated in solutions and/or
partitioned through precipitation or liquid phase separation. Some of these species
could have been sampled by the Cassini INMS and CDA instruments, especially
in association with salt-bearing grains.
Aqueous alteration of rocks starts with partial and unequal dissolution of solids
followed by precipitation of secondary phases. Water is mainly consumed in hydration and oxidation reactions. As in chondritic materials, oxidation could involve
formation of magnetite, Fe-bearing olivine and phyllosilicates from Fe-Ni metal,
tochilinite from troilite, phosphates from phosphides, and Ni-sulfides from Fe-Ni
metal. Hydration leads to formation of phyllosilicates (serpentine, cronstedtite,
saponite, chlorite) and tochilinite. Precipitation of carbonates partially consumes
inorganic carbon and and cations (Ca2+ , Mg2+ , Fe2+ , NH+
4 ) from solution. Fluid
composition and pH becomes strongly affected by the solubility of newly-formed
minerals. Evolution of fluid chemistry leads to an alkaline NaCl solution with
subordinate amounts of carbonate and bicarbonate ions, and K+ (Zolotov 2007,
2012). Although Fe, Ni, Mg, Ca, Si, Al, and S are abundant in the rocks, they are
minor constituents of solutions owing to low solubility of corresponding secondary
minerals (mainly phyllosilicates and sulfides). Geologically rapid water-rock reactions during differentiation could have produced a NaCl- and NH3 -bearing solution
layer above the core. The solution composition should not be much different if the
differentiation occulted without formation of a large global water layer.
34
Sascha Kempf et al.
1397
Subsequent exhaustion of heat sources caused freezing of fluids and precipitation of salts, especially around the rocky core. Methane and CO2 could be
trapped in clathrates in most of the new-formed ice (section 2.2.2). Although CO2
clathrate is more stable, conversion of CO2 to HCO3 − and CO23− ions in alkaline
fluids lowered CO2 fugacity and constrained clathrate formation.
The composition of plume gases indicates survival of some accreted icy material
during differentiation. This is consistent with the thermal evolution models of
Schubert et al (2007) that do not imply melting of a primordial icy-rocky shell of
accreted materials. However, the lack of detection of rocky particles by the Cassini
CDA instrument is not consistent with gas origins in remnants of undifferentiated
crust. In addition, the crater record (section 2.1.5) and the composition of the
Enceladus’ surface and plume solids do not indicate occurrence of rock-bearing
materials in the upper surface layers. These observations could be explained by
heterogeneous accretion (Charnoz et al 2011), a gravitational submergence of the
initial shell, and cryovolcanic activity within the primordial shell. In a case of
heterogeneous accretion, a low-density volatile-rich icy shell could have floated
above an early ocean. In a case of early homogeneous accretion, a strong heating
by 26 Al caused thinning (Schubert et al 2007) and submergence the rock-bearing
shell followed by its partial or complete melting.
Cryovolcanic activity could have been driven by disruptions in the primordial shell caused by impacts, tidal motions, and volume changes due to ocean
freezing (e.g., Manga and Wang (2007)) and hydration/dehydration processes in
the core. The non-volatile chemistry of rapidly frozen dikes, cryovolcanic flows
at the surface, and ancient plume deposits would represent the composition of
aqueous fluids. However, low-solubility gases (CO, N2 , H2 , CH4 and other hydrocarbons) could have been lost from dikes, cryovolcanic flows, and eruptive vents.
Decompression also leads to some CO2 degassing through shifting of the carbonate
equilibria (CO23− → HCO3 − → CO2 ) in alkaline NaCl-NaCO3 /NaHCO3 solutions
suggested in the interior (Postberg et al 2009; Zolotov 2007). In deep dikes and
other intrusions of aqueous fluids, CH4 and CO2 could be trapped in clathrates together with salts precipitated upon freezing within the primordial icy-rocky shell.
Other oceanic impurities (NH3 , methanol, and other organic species), organic liquid phases, and low-solubility gases (H2 , CO) could be trapped in gas/liquid inclusions in ice. Ices formed through cryovolcanism should be compositionally distinct
from accreted ices. It follows that the comet-like concentrations of CH4 and NH3
in plume gases are not consistent with gas origin in secondary ices.
1398
3.3.2 Chemical processes related to tidal heating
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The presence of salts, clathrate hydrates, ammonia and/or methanol in water ice
is known to affect the rheological properties of the mixture (Arakawa and Maeno
1994; Cole et al 1998; De La Chapelle et al 1999; Durham et al 2003, 2005, 2010;
McCarthy et al 2011). As a consequence, their presence in the icy shell is expected
to modify the response to the tidal forcing. At temperature near the melting point,
the presence of salts, ammonia, and methanol favor the formation of melt pockets at grain boundaries, owing to their anti-freezing properties (e.g. Fortes and
Choukroun (2010)). Experimental tests on water ice containing brine pockets (De
La Chapelle et al 1999) indicates that a melt fraction of about 5% decreases the
viscosity of ice at the melting point by about one order of magnitude (Tobie et al
2003). Cyclic loading experiments on terrestrial sea ice (Cole et al 1998) indicate
that the viscous strain rate increases markedly with brine volume, suggesting that
tidal dissipation should be significantly enhanced in regions enriched in brine pockets. Brine pockets could, for instance, be entrained by convective motions in the
icy shell and be concentrated in warm upwellings (Zolotov et al 2004; Pappalardo
and Barr 2004)).
Ammonia is expected to strongly impact the tidal dissipation process owing to
its very strong effect on the melting temperature. Laboratory experiments show,
indeed, that partially melted ice containing 4% of ammonia would have a viscosity ranging between 109 and 1011 Pa s at temperatures 180-210 K (Arakawa and
Enceladus as an active body
35
1448
Maeno 1994), whereas the viscosity of pure water ice at the same temperature
would be at least 1020 Pa s. A slush layer enriched in ammonia may form subsequently to the downward freezing of an internal ocean, since the concentration
of ammonia progressively increases in the remaining liquid. An isoviscous slush
layer of 1010 Pa s could, for instance, produce a total dissipation power of about
2 GW (Tobie et al 2008). Likewise, tidal dissipation is expected in a case of salty
brines and/or organic liquids accumulated at the core-icy shell interface (Zolotov
2007). Dikes of enriched inorganic and organic liquids that may be injected in
the icy shell during crystallization episodes (see section III.2.2) should also locally
enhance the tidal dissipation rate. This will help strike-slip motions between different icy blocks (e.g. Nimmo et al (2007)). Local tidal dissipation would cause
further melting which, in turn, induces dissipation. However, significant melting
would limit tidal dissipation. Further modeling is needed to understand the complex coupling between anti-freezing compounds, tidal dissipation, and convective
motions.
At temperatures far below the melting point of pure water ice, the presence of
ice impurities (clathrates, salts or silicate microparticles) usually tends to make
the ice mixture more resistant to creep (e.g. Durham et al (2003, 2005); McCarthy
et al (2011)). However, mixtures of water ice and hydrated salts appear to be more
brittle than pure ice (McCarthy et al 2011). This may favor the formation of a
zone of weakness promote tidal deformation in the icy shell. Even small amounts
of impurities also have an effect on the evolution of the ice grain size (e.g. Durand
et al (2006)). In polar ice sheets on Earth, for instance, soluble and second-phase
particles located at grain boundaries affect grain growth by exerting a pinning force
on moving grain boundaries (e.g. Alley et al (1986)). As the deformation of ice is
sensitive to grain size in the stress conditions expected within Enceladus’ ice shell
(Barr and McKinnon 2007), the impurities, even if present in small concentrations,
could have a major impact on the rheology of ice. Both experimental investigations
and numerical modeling are required to understand the interplay between solid
impurities, rheology, and tidal dissipation.
1449
3.3.3 Habitability of Enceladus
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The likely occurrence of aqueous fluids, organic species, and the detection of carbon
species with diverse oxidation states (CO2 , CH3 OH, CH4 ) imply habitable environments in the interior of Enceladus. Although the composition of salt-bearing
particles emitted from Enceladus (Postberg et al 2009, 2011) indicates present
or past aqueous environments it does not provide direct indication for life. The
composition of plume gases does not indicate their aqueous sources, though the
deeper interior could contain altered species transformed in abiotic (section 3.3.1)
or biologic processes. Likewise, the comet-like isotopic ratio of H in plume gases
(Waite et al 2009) may not characterize deep water-bearing environments.
If life exists, it may not be limited by the ability of C-N-O-H-bearing nutrients and biologically-important metals, though the low solubility of phosphates
in alkaline fluids may confine biological productivity. The moderately alkaline pH
evaluated for Enceladus’ fluids (Postberg et al 2009; Zolotov 2007) is suitable for
known microorganisms. However, organisms may not live in extremely cold and
alkaline NH3 -rich fluids (at 173 to ∼230 K) that are characterized by low activity
of water.
Putative life on Enceladus could be presented by chemotrophic microorganisms
that gain metabolic energy from oxidation-reduction reactions (McKay et al 2008;
Parkinson et al 2008). Major metabolic pathways could involve reduction of C
species (CO2 , CO, formaldehyde, methanol, and unsaturated hydrocarbons) by
H2 produced through water-rock interactions (McKay et al 2008). Methane is
the final product of these reactions. Aqueous hydrogenations of formaldehyde,
and unsaturated hydrocarbons to produce methanol are energetically favorable
as well. However, traces of O2 produced through radiolysis of ice or liquid water
(e.g., via radioactive decay of 40 K, Chyba and Hand (2001)) may not be used by
organisms that oxidize H2 , C, S, and Fe2+ species because of rapid consumption of
36
Sascha Kempf et al.
1489
O2 through oxidation of sulfides to native S and/or sulfates. The abiotic formation
of some sulfates creates a potential for biological sulfate reduction by H2 , CH4 , and
other organic species (McKay et al 2008). However, a large-scale abiotic oxidation
of sulfides may not have occurred in cold and/or short-living aqueous environments
with a limited recycling of surface oxidants. Therefore, the mass of sulfate available
for biological reduction could be negligible.
In the history of the moon, redox disequilibria could have been activated
through a melting of volatile-rich ices following by water-rock-organic interactions. Organisms may have adapted to changes of prevailing redox disequilibria
and evolving water activity. The organisms could have been able to survive (in a
dormant stage) in sub-freezing environments without aqueous fluids and in neareutectic brines rich in ammonia, chlorides, and organic liquids at temperatures
>160-170 K. Signs of life could be potentially detected in plume components that
indicate aqueous origins (e.g. salty ice grains).
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3.4 Summary
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The surface materials and plume emissions are mainly presented by water (crystalline and amorphous ice and vapor). Enceladus’ emissions also contain CO2 ,
NH3 , CH4 , CH3 OH, organic species with mass >100 Dalton, chlorides and carbonates/bicarbonates of Na and K, and nanometer-sized sillica grains.
Although the composition, size and dynamics of the emitted ice particles
strongly point at a source of liquid salty water, there is no unequivocal chemical
evidence for a large water reservoir that supply the plume components. Solution
pockets enriched in ammonia are stable in the warm water ice at the surface but
the composition of plume gases is not consistent with degassing of aqueous fluids
and implies a contribution from sublimation of heated ice that contains volatiles in
clathrates, brine pockets, or structural defects. The composition of emitted gas end
solids points at a diversity of processes creating the plume. A mixture of liquid and
solid sources is the most likely scenario. A south polar low-viscosity layer between
the rocky core and the ice shell is likely although it is unclear how it would connect
to possible saltwater pockets at shallow depth feeding the plumes. The possible
occurrence of primordial newer melted volatiles together with aqueously-processed
compounds in the icy shell suggests incomplete differentiation of the moon and
may indicate spatial separation of these diverse species. The compositional heterogeneity should reflect mechanisms of water-rock differentiation and subsequent
thermal processes in the icy shell. Chemical data indirectly indicate predominantly
low-temperature (≤273 K) water-rock interaction in Enceladus’ history. This implies a low amount of accreted 26 Al and a late, possibly heterogeneous accretion
is consistent with the current formation models for the Saturn system.
Although Enceladus could be habitable and may harbor chemotropic organisms
(e.g., methanogens) in subsurface aqueous environments, current compositional
data do not indicate biological processes.
1517
4 Plume formation and properties
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The Enceladus activity is signified by the plume towering the south pole of the
satellite. The plume allowed for a direct in situ sampling of gas and particles during Cassini flybys, providing unique information on the composition of material
stemming from the interior of the moon. Moreover, the physical characteristics of
the plume pose important constraints on the activity in general, such as the production rates of vapor and particles and their variability, the ejection speeds, the
particle size distribution, and the distribution of plume sources and their relation
to features and temperatures on the surface.
Enceladus as an active body
37
Fig. 13 Left panel: Cassini image PIA08386 (NASA/JPL, http://www.ciclops.org), showing
a high phase image of the Enceladus plume. Visible are numerous dust jets but also a broad distribution of dust. Right panel: Cassini image PIA11688 (NASA/JPL, http://www.ciclops.org).
A close-up view of the south polar terrain at large phase angle, bringing out details of the dust
ejection. In this particular viewing geometry the terminator crosses the Tiger Stripes. Dust
emerging from those sources on the dark side of the shadow becomes visible at hight when
leaving the shadow. Arrows mark the shadow boundary, which becomes visible because of the
presence of dust.
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4.1 The gas and dust components of the plume
The dust plume is directly seen in the high phase images from the Cassini ISS
(Porco et al 2006; Spitale and Porco 2007; Ingersoll and Ewald 2011), in the near
infrared data from Cassini VIMS (Hedman et al 2009; Hedman et al 2013) and it
has been recorded in situ by the Cassini CDA (Spahn et al 2006b; Schmidt et al
2008; Kempf et al 2010; Postberg et al 2011). This part of the dust configuration
is formed by grains composed dominantly of water ice (Postberg et al 2011) and
sized in the sub-micron to micron range (Spahn et al 2006b; Hedman et al 2009;
Postberg et al 2011). Additionally, jets of very fine, charged particles were detected
by the Cassini CAPS instrument (Jones et al 2009; Hill et al 2012). The properties
of these jets are discussed in detail in section 5.2.3 and we focus in this section
primarily on the larger grains.
The sources of the dust plume are the Tiger Stripes, with a spatially variable
level of activity. Several prominent sources were identified with particularly hot
regions on the Tiger Stripes (Spitale and Porco 2007). A later, preliminary analysis
of images reports a much larger set of sources (Porco et al 2011). The distribution
of gas sources is perhaps best constrained from stellar and solar occultations by
the plume recorded with the Cassini UVIS experiment . In a similar manner as
for the dust plume, the UVIS data shows evidence for a broad ejection of gas with
additional peaks in gas density, which were shown to be correlated (Hansen et al
2008, 2011) with the most prominent dust jets (Spitale and Porco 2007).
The gas and the dust phases of the south-polar ejection exhibit important
differences, though, in spite of the fact that they likely originate from the same
sources. Generally, the gas ejection speeds are found to be on the order of several
times the escape speed (≈ 237m/s) (Hansen et al 2006; Tian et al 2007; Smith
et al 2010; Dong et al 2011) while the scale hight of the dust plume suggests mean
ejection velocities on the order of 100 m/s only (Porco et al 2006; Schmidt et al
2008; Hedman et al 2009). This dynamical difference has important implications.
Namely, it shows that there cannot exist a strong momentum transfer from the dust
to the gas, simply because the gas would then be efficiently slowed down by the
particles, which is not seen. On the other hand, for plausible subsurface gas densities
(Schmidt et al 2008) one expects a strong momentum transfer from the gas to the
dust, mediated by friction (Epstein 1924; Cunningham 1910). One possibility is
that particles are decelerated relative to the gas flow already in the subsurface
vents. This has been explored in detail by Schmidt et al (2008), who suggest that
particles, entrained in the subsurface gas stream are repeatedly decelerated by
collisions with the walls of the vents. Another principal possibility is that both gas
38
Sascha Kempf et al.
1567
and dust emerge at slow, subsonic speed from the vents and the gas is accelerated
further by pressure forces, when expanding above the vents, and before the gas
becomes collisionless. The latter scenario has not been explored to date and it is
not clear if this is realistic for the observed gas and dust densities, as well as the
distribution of velocities and grain sizes.
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4.2 Constraints
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The observed properties of the gas and dust plumes pose important constraints
on possible scenarios of plume formation. In this section we summarize such constraints and discuss models suggested in the literature in this light.
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4.2.1 Steadiness and variability
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In data from stellar and solar occultations by the plume observed by UVIS there
is little evidence for a variability of the vapor production rate (Hansen et al 2011)
within the error limits of the data reduction. Especially, no variation was seen in
correlation with the orbital phase of Enceladus, which was predicted as a consequence of the tidal flexing of the south polar terrain and the Tiger Stripes (Hurford
et al 2007), although the coverage of orbital longitudes with occultations was limited. Saur et al (2008) modeled the influence of the water plume on the corotational
plasma and derived the corresponding magnetic field perturbation. From comparison to data obtained by the Cassini MAG instrument at multiple flybys they infer
plume source rates that exhibit a significant level of variability (see Tab. 3). Dong
et al (2011) fitted a model of eight water vapor sources, identified with the jet
sources by Spitale and Porco (2007), to measurements by the INMS instrument
recorded during the E3, E5, and E7 encounters. They found ejection speeds and
plume densities consistent with those inferred by UVIS. In contrast, Smith et al
(2010) infer a variability by a factor of four between E3 and E5 (seven months).
In a recent analysis of VIMS data Hedman et al (2013) find a periodic variation of the dust plumes’ brightness that correlates well with the orbital phase
of Enceladus (i.e. with a period of ≈30 hours). This supports well earlier ideas
on the influence of tidally induced normal stress on the width of the Tiger Stripe
cracks (Hurford et al 2007), possibly with a phase shift between forcing and response due to the effect of a physical libration of Enceladus (Hurford et al 2009).
Such a tidally controlled periodicity of the Enceladus activity has consequences
for the interpretation of data and modeling, which need to be explored in future
investigation. The observed fast modulation of the plume brightness is primarily a
consequence of the variation of the cross section of the vents over the tidal cycle,
leading to a corresponding variation of the gas and dust fluxes. It is not expected
that the underlying subsurface distribution of heat varies noticeably on the same
timescale.
Besides this periodic brightness variation with the tidal cycle the plume appears
to be continuously active at least over the time span of the Cassini observations
and the question arises how long has it been active in the past. One lower limit on
the duration of the activity comes from the lifetimes of the E ring particles, because
the plume must sustain this ring. These lifetimes vary strongly with particle size
(Kempf et al 2010), but it seems that the directed, vertical injection to the E
ring of particles through the south polar plume tends to increase the lifetimes of
particles (Juhász et al 2007) when compared to pre-Cassini models of the ring
(Juhász and Horányi 2002; Juhász 2004). Overall, it should take roughly 100 years
to build up a steady E ring. But most probably the timescales for plume activity
rather correspond to the much longer geological timescales, comparable to the ages
inferred for the south polar terrain itself, which is on the order of 107 years or less
(Porco et al 2006).
One age constraint was derived from certain features on the Enceladus surface
seen in images, which were interpreted as deposits of plume particles (Schenk et al
2011). The comparison of the estimated deposit thickness to accumulation rates
Enceladus as an active body
39
1621
obtained for the same location from ballistic particle models (Kempf et al 2010)
yielded a duration of 10s of millions of years for the plume activity. However, this
preliminary analysis needs to be extended to a larger set of images, from multiple
locations on the surface, so that the effect can be unambiguously attributed to the
action of plume particles falling back onto the surface.
1622
4.2.2 Dust to gas mass ratio
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One important constraint for models of plume formation is the dust to gas ratio.
The total dust mass in the plume was derived by Ingersoll and Ewald (2011)
from photometry of the plume and an ice to vapor ratio of 35 to 70%. At least
the upper limit appears uncomfortably large for pure condensation models, even
if one takes into account the afore mentioned effects. For instance, from models
invoking condensation as the process of particle creation one finds that typically
about 5 to maximally 10% of the gas discharge can be condensed (Schmidt et al
2008), depending on the initial enthalpy of the gas and the level of supersaturation
reached in the flow. On the other hand, from these numbers follow significantly
larger dust to gas ratios in the plume along a given line of sight. In practize, such
line of sight densities are inferred for the gas from occultations and from the local
brightness of the dust plume. When comparing to the bare source rates one needs
to take into account that the dust is ejected at slower speeds, thus contributing to
the density in a given volume element of the plume more efficiently than the gas.
Also, the gas is faster than the escape speed, and thus, each molecule traverses
the plume only once before it escapes the satellite. In contrast, most of the dust
grains fall back, so that they traverse the plume twice before they re-impact the
surface.
An additional contribution to the plume mass may come from particles that are
entrained in the gas flow as frozen droplets from a spray above a liquid reservoir,
as it was proposed by Postberg et al (2009, 2011). In this case only a part of the
particulate mass would condense. The details of this model (energy budget, size
distribution of grains) have not been worked out to date, but the concept might
be consistent with the model proposed by Matson et al (2012), to explain the
subsurface transport of heat and material on Enceladus by a cycle driven by the
exsolution of gases. Bubbles produced by such gases could generate spray above
the liquid, with droplets that have a composition identical to the liquid.
1650
4.2.3 Grain size and velocity distribution
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Particle sizes in the plume were derived in situ by Cassini CDA (Spahn et al
2006b). For the flyby E2, which traversed relatively distant parts of the plume,
they inferred a size slope of ≈ −3 for the differential size distribution. In the E
ring radially close to Enceladus (Kempf et al 2008) inferred steeper differential
size slopes between −4 and −5, indicating that smaller particles have a higher
probability to escape the moon’s gravity field. From dust impacts on the Cassini
RPWS instrument a yet steeper distribution with a power law exponent around −6
was inferred (Kurth et al 2006). The likely explanation for the different expontents
is that the true distribuiton of grains cannot be approximated well as a power law
in the full range of particle sizes covered by various instruments. Probably the
distribution falls of more rapidly than a power law and CDA fits exponents to a
flatter part of the distribution for sizes from sub-micron to about two microns,
whereas RPWS sees the steeper distribution in the few micron range.
For the velocity distribution of particles leaving the Enceladus vents Schmidt
et al (2008) derived the form
fR (u) du =
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R
Rc
1+
R
Rc
u
u2gas
1−
u
ugas
1− RR
c
du,
(3)
where fR (u) is the differential velocity distribution that depends parametrically
on the particle radius R. The critical radius Rc lies in the submicron range and it
40
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Sascha Kempf et al.
depdends on the dynamical coupling between gas and grains. Here, ugas is the gas
velocity. A comparison to VIMS data (Hedman et al 2009) suggested Rc ≈ 0.4µm.
Formula (3) follows from the statistics of particles that are decelarated repeatedly
by collisions with the walls of the vents and then are re-accelerated by gas friction.
It can, however, more generally be regarded as a convenient parameterization for
the size-dependent speed distribution. The distribution is normalized to unity in
0 ≤ u ≤ ugas . Then formula (3) implies that grains with sizes R > Rc have a speed
distribution that peaks at a velocity between 0 and ugas , while grains with sizes
R < Rc have a distribution with one single pole at u = ugas .
Using equation (3) one can construct the combined distribution for particle
sizes and ejection speeds
f (R, u) = fR (u) p(R)
(4)
1696
by multiplication with a suitable size distribution p(R). For instance, Postberg
et al (2011) used a power law for p(R) to construct from equation (4) a plume
model with a size distribution that varied with position in the plume (see also
section 3.1.3 and figure 7).
From the analysis of VIMS images taken from the Enceladus plume at high
phase angles, with a high spectral resolution in the near infrared, Hedman et al
(2009) reconstructed the size distribution in the plume and its variation with
altitude above the south pole. They find that generally the sub-populations of
larger particles have smaller scale hights than the small ones. The size distributions
at lower altitudes are found to be shallower than a power law with exponent −
in the range from 1 to 2 microns but are steeper than such a power law at larger
size. Generally, a fairly steep drop is observed in the particle abundance for sizes
larger than 3 micron. Hedman et al (2009) assumed that the dust density in
their measurement is dominated by those grains reaching their maximal altitude
where their line of sight interstects the plume. Using the integrals from the two
body probelm they could then invert their data to reconstruct a distribution of
starting velocities. The size-dependence of the slopes of the so determined velocity
distribution is roughly consistent with formula (3).
1697
4.2.4 Supply rate of particles to the E ring
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Juhász and Horányi (2002) derived an E ring model with particles generated at
various satellites in the E ring. The evolution of the particles was subject to Saturn’s gravity (including higher order contributions due to the oblate shape), solar
radiation pressure, and electrodynamic forces, and plasma drag, with an appropriate charging model to obtain the instantaneous grain charge in the magnetosphere.
The sinks of the particles are collisions with Saturn’s A ring, the E ring moons
and plasma sputtering of the grains. This model did not take into account the
directed injection of particles from Enceladus through the south polar plume. A
steady state E ring was obtained at a production rate of about 1 kg/s particles at
Enceladus.
Spahn et al (2006b) fitted a model of particles ejected from Enceladus’ south
polar terrain to CDA data from the flyby E2. Their best match with the model
suggested a supply rate of 0.2 to 1 kg/s from Enceladus to the ring. Porco et al
(2006) derived from the analysis of Cassini images an escape rate of 0.04 kg/s.
Schmidt et al (2008) constructed a model plume based on equation (4) and particle size-distributions obtained condensation in the subsurface vents. Parameters
of the model were constrained from fits to CDA data from E2 and to brightness
profiles of images. They obtrained a mass production rate of 5kg/s from which
0.5 kg/s escape to the E ring. From high phase images of the plume Ingersoll
and Ewald (2011) derived an about 10-fold value for both the production and the
escape rate.
1719
4.2.5 Composition: Sodium and volatiles
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The composition of the water plume has been derived from INMS (Waite et al
2006, 2009, 2011) and UVIS data (Hansen et al 2006, 2008, 2011). The current
Enceladus as an active body
41
1731
interpretation of these data is summarized in section 3.1.2 and the implication for
liquid and solid sources of the gas are discussed in section 3.2.4. The ice grains in
the E ring and in the plume (Postberg et al 2009, 2011) were found to split into
different compositional types (see section 3.1.3). Importantly, most of the grains
show at least small amounts of sodium which can be attributed unambiguously to
the presence of salt. Moreover, (Postberg et al 2011) have shown that most of the
ice mass in the dust plume has a salinity on the order of 1%. On the other hand,
from earth-bound observation Schneider et al (2009) showed that the gas in the
plume is practically free of sodium, from analysis of the brightness of the plume
at the wavelength of the sodium D2 line.
1732
4.3 Scenarios for plume formation
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4.3.1 Explosive boiling model
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This scenario for plume formation was proposed by (Porco et al 2006). It assumes
that liquid is present very close to the surface (motivated by the high surface temperatures seen by CIRS (Spencer et al 2006)) and that fractures, rapidly opening
in the ice, suddenly expose the liquid water to vacuum. In this zero pressure
condition the uppermost layer of liquid should instantaneously begin to boil while
simultaneously part of it begins to freeze. Part of the frozen phase would be ejected
as plume particles, accelerated and dragged to space by the expanding gas. This
process occurs far from equilibrium, and thus, in principle large solid to gas ratios
should be possible. But the process is also self-limiting, because any crack should
be rapidly sealed by the freezing part of the liquid. This is hard to reconcile with
the ongoing quasi-continuous activity, as is observed, even if one invokes multiple
cracks to maintain an average level of activity. Moreover, in this fast process the
gas from the boiling phase, as well as the frozen part should have the same sodium
content as the liquid from which they originate. But this is not observed: the plume
vapor is essentially sodium free (Schneider et al 2009) compared to the 1% salinity of the particles (Postberg et al 2011). Brilliantov et al (2008) have estimated
for the velocity of the expanding gas from basic thermodynamic principles. They
found that the maximum speeds should be as low as meters per second, which is
not consistent with observations, implying gas velocities on the order of several
hundred meters per second or higher (Hansen et al 2008).
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4.3.2 The clathrate reservoir hypothesis
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Kieffer et al (2006) proposed a dry model for the plume activity that is based on
the decomposition of gas hydrates (clathrates) in a subsurface reservoir (see also
Gioia et al (2007b)). This is motivated by the observation that plume gases like
CH4 and CO2 can form clathrates, and then exhibit higher solubility in water
ice than in liquid. These clathrates become unstable at high temperatures and/or
low pressures so that near cracks, or heat sources, they would decompose into
(volatile) gas and solid ice. If there exists a shell of clathrate-rich ice below the
south-polar terrain, a decomposition front could propagate from the Tiger Stripes
into this reservoir. The advection of the gases to the surface through the Tiger
Stripe faults would explain the volatile gases found in the plume. On the other
hand, the process does not produce water gas directly, which would need to form
under subsurface conditions by sublimation from the icy break-up products of the
clathrates. This might be difficult to achieve thermodynamically, especially at the
relatively low temperatures for which the model, also called “frigid faithful” was
proposed. The pressure of the gas should rapidly adjust to the saturation vapor
pressure, which is a strongly dropping function of temperature. The most severe
difficulty of this model, as for any dry scenario, is that it cannot account for the
observed sodium in the plume grains (Postberg et al 2011).
42
Sascha Kempf et al.
Fig. 14 Steady state temperature profiles in a vertical plane cutting a shear zone in the
south polar ice cap. The original model was proposed by Nimmo et al (2007) and the figure is
reproduced from Ingersoll and Pankine (2010).
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4.3.3 Shear heating models and sublimation from warm ice
1802
The production of heat by tidally induced shearing motion in the south polar
ice cap was modelled by Nimmo et al (2007). The frictional heating results from
strike-slip motion along faults (the Tiger Stripes) in the upper, brittle ice layer and
from viscous heating in the ductile ice at depth. The mechanism is efficient if there
exists an ocean between the rocky core of the satellite and the ice mantle, so that
the tidal displacements of the latter can be large. The heat produced in the ice
can be transported by conduction, by the hydrodynamic flow of vapor through the
cracks, and by diffusion of vapor if the ice is porous. Nimmo et al (2007) considered
a balance of non-equilibrium vapor production and diffusion while Ingersoll and
Pankine (2010) suggested to limit the vapor density so that it locally adjusts to
the saturation vapor pressure (see also discussion in section (2.2.4)).
As a result, one expects the ice to warm up along a relatively thin zone that
extends from the Tiger Stripes downward into the ice cap. The process reaches
an equilibrium temperature distribution as depicted in figure (14), resulting from
energy input in the shear zone, redistribution of heat by conduction, vapor diffusion
and advection, and loss due to radiation at the surface and vapor escaping to
vacuum. Additionally, heat may be effectively transported in form of latent heat
of the vapor. It can be transferred to the ice when vapor re-condenses on the walls
in the upper part of the faults. In total, on the surface this would result in a heat
flow pattern that is localized at the Tiger Stripes (see Fig.3 c).
This process is capable of producing abundantly water vapor, possibly reaching the measured mass fluxes, if the temperatures reach sufficiently high values,
depending on the shear rate, the unknown thickness of the ice and poorly known
parameters, as the viscosity of the ice. But the model cannot account directly for
the observed non-water gases. To this end, an additional process must be invoked,
like outgasing of volatiles contained in the ice, similar to the clathrate decomposition model. The particles of the dust plume might be created in this scenario
by condensation in the vents as proposed by Schmidt et al (2008). However, the
model cannot account for the presence of sodium salts in the ice particles.
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4.3.4 Evaporation from liquid
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Conditions for melting in the shear heating model were investigated by Ingersoll
and Pankine (2010). One main conclusion from that work was that melting can
indeed occur even if the heating process is self-limiting because of the temperature dependence of the viscosity of the ice. However, the precise conditions would
Enceladus as an active body
43
Fig. 15 Figure reproduced from Postberg et al (2009). Formation of salt rich droplets by dispersion from the liquid (a) which subsequently freeze and grow to some extent by condensation
of vapor. Salt poor grains could arise directly from homogeneous nucleation in the vents (b).
1834
depend on the unknown grain size in the ice (because this determines the porosity
and thus the magnitude of vapor diffusion) and the width of the cracks (because
this would limit the amount of heat advected with the hydrodynamic vapor flow).
If the melting affects brine pockets in the subsurface ice, then there is no principal
problem with having sodium in the grains. Such frozen brine pockets may result
from previous phases of activity in the Enceladus history in a scenario of an intermittent activity cycle (Roberts and Nimmo 2008), and they might be part of a
convection system in the ice shell (Tobie et al 2008) (see also Fig. 2).
The liquid scenario offers a simple explanation for the detection of sodium salts
in the ice grains, if the liquid reservoir is, or was in contact with the rocky core
Enceladus, enriching the water with minerals. In fact, the most prominent species
predicted for an Enceladus ocean (Zolotov 2007) were identified in E ring and
plume grains (Postberg et al 2009, 2011) by Cassini CDA. These CDA compositional measurements are described in detail in section (3.1.3). The interpretation
of the results is shown in figure (15), suggesting that salt rich grains are formed
from droplets formed by dispersion from a salty liquid. These droplets would be
sustained in the vapor above the liquid and become entrained in the quasi steady
flow of gas to the surface through back-pressurized vents.
One possible mechanism for dispersion of droplets from the water could be
bubbles of exsolved gases (Matson et al 2012). The droplets would freeze and be
transported to the surface by the nearly steady gas stream through the vents. The
salt-poor grains, and especially the nano-grains seen by CAPS (Jones et al 2009;
Hill et al 2012), could form by homogeneous nucleation from the supersaturated
water vapor. In such a scenario the water gas could be formed by evaporation from
both the liquid and the surrounding warm ice (Fig. 16). Additional volatile gases
could be contributed from the ice, possibly created by clathrate decomposition
close to the warm shear-zone and transported by diffusion.
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5 Interaction with Saturnian system
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5.1 Introduction
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Enceladus and its plume do not exist in isolation, but are under the influence of
their immediate environments. The moon and plume are constantly bombarded
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Fig. 16 Schematic scenario for the formation of plume gases and dust.
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by magnetospheric thermal plasma and the energetic particles of Saturn’s radiation belts. This bombardment liberates material through sputtering, affecting the
chemistry of the plume. Strong currents, set up between the ionized plume and
Saturn itself, provide energy input into the planet’s upper atmosphere. In return,
Enceladus is the primary source of material for the E ring, and is also the ultimate source of material contained in the neutral and ionized tori that encircle
the planet. These tori exert a strong influence on the nature of the planet’s entire
magnetosphere.
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Here, we review our current understanding of this complex set of interactions. A
schematic overview of the domains directly or indirectly influenced by the moon’s
activity is given in Fig. 17.
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5.1.1 Discovery of Enceladus’s magnetospheric interaction
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Enceladus’s unusual nature as an extremely high albedo object, as revealed by the
Voyager probes, made encounters with it a high priority when planning CassiniHuygens’s tour of the Saturn system. As described below, its coincidence with the
densest region of the E ring identified it as the likely source of particles within
that ring (Baum et al 1981). Shemansky et al (1993) reported the presence of
an OH torus, again seemingly associated with Enceladus’s orbit. Evidence of a
magnetospheric interaction was not as forthcoming. Pioneer 11 and Voyager 1
and 2 data revealed the presence of depletions in energetic particle fluxes termed
micro- and macro-signatures that are also present at other moons’ L shells (van
Allen et al 1980; Carbary et al 1983).
It fell to Cassini to reveal the extensive nature of this perplexing moon’s interaction with the Saturn magnetosphere. Table 1 lists Cassini’s encounters with
Enceladus to date. Encounter E0, in February 2005, was a relatively distant encounter, but during that approach to the moon, the magnetometer (Dougherty
et al 2004) detected a significant deviation in magnetic field strength and direction that indicated that the Enceladus-magnetosphere interaction was a strong
one, that the scale of the interaction was larger than the moon itself, and that the
moon possessed an ionized atmosphere (Dougherty et al 2006). As a result, the E2
encounter was lowered, and the data from the magnetometer and other particles
and fields instruments began to reveal the true extent of the moon-magnetosphere
interaction (Kivelson (2006) and references therein).
Enceladus as an active body
45
Fig. 17 Schematic diagram showing particles and plasma originating at Enceladus and their
interactions within Saturn’s magnetosphere. Only the major plasma processes (yellow ovals)
and affected particle populations (pink rectangles) are shown. While the single particle source is
Enceladus cryo-volcanism and surface outgassing, the ultimate particle sinks (blue rectangles)
are interplanetary space and the solar wind in addition to the atmospheres of Saturn and
Titan.
Parameter
Magnetic field B0
Plasma number density n0
Plasma velocity u0
Average ion mass m
Electron beta βe
Ion beta βi
Ion gyroradius rg
Alfvnic Mach number MA
Sonic Mach number MS
Magnetosonic Mach number MMS
Values
325 nT
70.5 cm-3
26.4 km s−1
17.6 u
0.0004
0.0094
14.82 km
0.131
1.32
0.03
Table 2 Typical plasma parameters observed in the vicinity of Enceladus (after Kriegel et al
(2009) and references therein)
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5.1.2 Cassini’s encounters with Enceladus
Before providing an overview of Cassini’s results, it should be borne in mind that
our exploration of the moon-magnetosphere interaction region is far from complete.
As tidally-locked Enceladus resides in a continuous flow of magnetospheric plasma
almost comoving with Saturn’s rotation, one hemisphere is constantly bombarded
by thermal plasma. As the plume source surrounds the moon’s southern pole, a
north-south asymmetry in the distribution of neutral and ionized gas, as well as
particulate matter, also exists. Cassini’s trajectory during each encounter with the
moon is essentially a straight line on the scale of the interaction region, therefore
to fully understand the interaction, many encounters are required. As well as only
having limited coverage in terms of encounter geometries, i.e. up- and downstream,
north and south, of the moon, as the spacecraft is three-axis stabilized with most
instruments having limited fields of view which have tight pointing constraints, we
have not yet obtained all observations required to put together a complete picture
of the interactions occurring at the moon. In no particular order, the instruments
whose results constitute the majority of the material presented here are the Magnetospheric Imaging Instrument, MIMI (Krimigis et al 2004), the Magnetometer,
46
Sascha Kempf et al.
Fig. 18 Cassini encounter geometries, separated according to the four types of encounter.
Top panels: view looking along the thermal plasma wake, north upwards, corotation flow out
of the page, Saturn towards the right. Bottom panels: view from the north; corotation from
from top, Saturn towards the right. [The figure pairs need to be boxed to together to make
the association more clear]
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MAG (Dougherty et al 2004), the Cassini Plasma Spectrometer, CAPS (Young
et al 2004), the Cosmic Dust Analyzer, CDA (Srama et al 2004), the Radio and
Plasma Wave investigation (Gurnett et al 2004), the Ion and Neutral Mass Spectrometer, INMS (Waite et al 2004), and the Ultraviolet Imaging Spectrometer,
UVIS (Esposito et al 2004). For the purposes of investigating the interaction with
the magnetosphere, Enceladus encounters can be classified by flyby geometry into
four groups, as follows:
Upstream encounters (E0, E1, E2, E10): Cassini passes upstream of the moon.
It does not encounter the thermal plasma wake, and does not pass directly
through the plume.
North-South plume crossings (E3, E4, E5, E6): Cassini’s velocity is primarily northsouth in the Enceladus frame. The spacecraft approaches the moon from downstream, then passes through the plume, emerging upstream. E4, E5, and E6
were very similar in geometry.
Southern encounters (E7, E8, E9, E11): Cassini passes directly through the plume
with only a small velocity component in the north-south direction. It remains
south of the moon throughout the encounter.
Northern encounters (E12, E13): Cassini passes directly north of Enceladus with
only a small north-south component of motion.
Enceladus as an active body
Encounter
E0
E1
E2
E3
Neutral source rate
1600 kg/s
(Saur et al 2008)
200 kg/s
(Saur et al 2008)
18 (or 36 or 72) kg/s
(Smith et al 2010)
200 kg/s
(Saur et al 2008)
300 kg/s
(Burger et al 2007)
E2 6kg/s
190 kg/s
(Smith et al 2010)
370 kg/s
(Dong and Hill 2009)
E4
E5
E6
E7
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Ion source rate
0.2 kg/s
(Saur et al 2008)
100 kg/s over 30 RE pickup region
(Pontius and Hill 2006)
0.2 kg/s
(Saur et al 2008)
3 kgs
(Burger et al 2007)
1- 3 kg/s over 5 RE pickup region
(Khurana et al.)
1- 3 kg/s over 5 RE pickup region
(Khurana et al.)
1- 3 kg/s over 5 RE pickup region
(Khurana et al.)
750 kg/s
(Smith et al 2010)
700 kg/s
(Dong and Hill 2009)
500 kg/s
(Dong and Hill 2009)
N/A
6.5 kg/s lost from neutral torus
(Leisner et al 2006)
Table 3 Estimates of neutral and ionized gas production rates at Enceladus.
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5.2 The Plume
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5.2.1 Production rate and composition
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The processes leading to the generation of the Enceladus plume are described elsewhere in this publication. By injecting significant quantities of neutral gases and
solid icy grains into Saturn’s magnetosphere, Enceladus acts as a major controller
of magnetospheric composition and dynamics (Kivelson 2006). The physical and
chemical processes occurring in the plume are numerous.
Estimates of the production rate of neutrals are varied, but most derived values
are of the order of hundreds of kg per second. The variability in the quantitative
estimates may largely be due to a real variation in the plume activity level; a
summary of production rate estimates is provided in Table 3.
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5.2.2 Ionization in the Torus and the Plume
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Several potential sources of plasma exist in the Enceladus torus and/or in the
Enceladus plume, including:
1. photoionization of neutrals by solar extreme ultraviolet (EUV) and x-ray radiation
2. impact ionization of neutrals by energetic magnetospheric ions and electrons
3. photoionization and impact ionization of E-ring grains and satellite (i.e., Enceladus) surfaces.
The main neutral species in the plume, as well as in the dense, main neutral torus
located at radial distances close to 4 RS , is water (H2 O), but water dissociation
products (i.e., OH, O, H) are also present (Waite et al (2006, 2009, 2011); Teolis
et al (2010)), particularly in the more extended outer torus (Johnson et al (2006);
Perry et al (2010); Jurac (2001); Richardson et al (1998); Melin et al (2009)).
Many other minor neutral species were detected in the plume by the INMS (Waite
et al 2009) including CO2 , C2 H2 .(See section xxxx of this review). The primary
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Sascha Kempf et al.
ionization reaction can be represented by:
hν + H2 O →H2 O+ + e
OH + + H + e
O+ + OH + e . . .
(5)
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The photoionization frequency due to absorption of solar radiation by water, is
very close to a value of 9 × 10−9 s−1 at the heliocentric distance of Saturn and
for solar minimum conditions (c.f., Burger et al (2007); Cravens, T. E., Ozak,
M. S., Campbell, I. P., Perry, M., Rymer A. M. (2011)). Secondary ionization by
photoelectrons contributes another 10 %, giving a total photoionization frequency
of 10−8 s−1 . The ion production rate in the torus and/or plume for low opacity
conditions (which is almost everywhere outside a distance of 1 RE ) is this ionization frequency multiplied by the neutral number density. In the core neutral
torus, photoionization appears to be the dominant ionization mechanism. Most
of the suprathermal electron distribution measured by CAPS ELS (Schippers, P.,
Andre, N., Johnson, R. E., Blanc, M., (Dandouras), I., Coates, A. J., Krimigis,
S. M., Young, D. T. 2009) appears to be explainable by photoelectrons associated
with photoionization, at least for electron energies less than about 70 eV. Measured
electron distributions and model distributions agree well for r ∼4-5 RS . The distribution appears to consist of a Maxwellian core ( ne ∼ 80 cm−3 and kTe ∼ 2 eV),
a photoelectron region (20 - 70 eV), and a more energetic tail region (E >70 eV)
(Schippers, P., Andre, N., Johnson, R. E., Blanc, M., (Dandouras), I., Coates,
A. J., Krimigis, S. M., Young, D. T. (2009); Cravens, T. E., Ozak, M. S., Campbell, I. P., Perry, M., Rymer A. M. (2011); Gustafson, G., Wahlund, J. E. (2010);
Sittler et al (2008)). At least near RS , the higher energy suprathermal tail exhibits
low enough fluxes that it should not contribute significantly to the total ionization
rate in the torus. However, for increasing radial distance from Saturn (r > 8 RS
or so), the more energetic electron population, as observed by the Cassini CAPS
ELS and MIMI experiments, appears to be more robust and is probably associated
with injection events from further out in the magnetosphere (Rymer et al 2008).
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5.2.3 Charged dust grains
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During close Enceladus encounters E3 (March 12, 2008), E5 (October 9, 2008),
E7 (November 2, 2009), E10 (May 18, 2010), and E13 (December 21, 2011), the
CAPS sensors’ apertures were oriented to encompass the ram direction, allowing
direct sampling of plume material moving slowly relative to Enceladus. During
the first of these encounters, IMS, ELS, and IBS unexpectedly recorded extremely
high ≥ 1 keV/q count rates within the plume. The IMS and ELS observations were
reported upon by Jones et al (2009). Populations of high energy magnetospheric
ions and electrons were quickly ruled out as the cause of these signatures as they
were only detected by CAPS anodes nearest the local ram direction.
The signatures were therefore almost certainly due to massive ions in the Enceladus plume. To estimate their masses, as is the case at Titan, as ions enter CAPS
sensors at known velocities, their measured kinetic energy per charge can be converted to a mass-to-charge ratio (M/q). The E3 encounter, with its encounter
speed of 14.4 km s−1 , turned out to have recorded negative particles with M/q
values of up to 26 600 u/q, and up to 46 500 u/q for positive grains, implying that
these particles are charged nanodust. As it is the relative speed of spacecraft
and plume material that determines the conversion factor between E/q and M/q,
lower speeds allow a greater M/q range to be covered by each CAPS sensor, at
the expense of M/q resolution. The encounters with suitable CAPS pointing involved near-constant encounter speeds ranging from 6.2 to 17.7 km s−1 . Encounter
E10, with its 6.2 km s−1 speed potentially allowed the detection of negative grains
with masses up to 144 600 u/q, compared to the fastest encounter E5, where its
17.7 km s−1 speed allowed only masses per charge of up to only 17 700 u/q to be
detected.
Enceladus as an active body
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Data from the E3 and E5 encounters remain the most striking obtained to
date, showing very high fluxes up to the top of the sensors’ E/q ranges (Jones
et al 2009); at lower energies, low mass positive and negative ions originating in
the plume gas itself were also detected (Tokar et al 2009; Coates et al 2010a). The
≥ 500 eV/ e0 signatures during these encounters imply that ELS detected negative
ions of M/q from 400 up to 26 600 and 17 600 u/q, respectively, and IMS detected
positive ions up to 46 500 and 30 700 u/q. The grains’ charge state is undetermined,
but if they are singly-charged, they measure up to > 2 nm radius, i.e., orders of
magnitude smaller than plume grains detected by other instrumentation (Spahn
et al 2006b). The derivation of the number density of the nanodust is currently
ongoing; a direct calculation of this parameter from the raw data is not possible
due to the many instrumental effects that need to be taken into account, including,
for example, the effects of ELS microchannel plate detection efficiencies for high
mass particles.
The source of these grains is the condensation of water vapour that emerges
at Enceladus’s south polar region at surface fractures. Modelling of this process
(Schmidt et al 2008) predicts the formation of grains on scales from nm to tens
of mm, with the smallest attaining velocities close to the gas speed.
Although some differences between signatures observed by the different elements of CAPS are due to the sensors’ time resolution, there is a clear mismatch
between the timing of fine scale structures observed in the positive and negative
grain data. In E3 IMS data, one clear, relatively broad flux peak is observed, preceded by a minor isolated peak. In ELS, two main peaks are seen: one detected
over ELS’s entire energy range and all anodes, possibly resulting from grain fragmentation inside ELS itself, and a second, more complex signature where the peak
energy shifts over time, indicating negative grains’ spatial separation.
The temporal, and therefore spatial differences between positive and negative
fine-scale structures are believed to be due to the effects of electromagnetic forces
on the nanodust. Initially, the motional electric field associated with the flow of
magnetospheric plasma past Enceladus will accelerate negative grains towards Saturn, and positive grains away from the planet. As reported by Jones et al (2009),
at a qualitative level, this appears to agree with the jet source locations catalogued by Spitale and Porco (2007). Detailed modeling of the processes moulding
the shape of the jets is ongoing (Arridge et al 2010), but it is clear that the electromagnetic fields present in the vicinity of Enceladus essentially allow the plume
region to act as a mass spectrometer, splitting the particles by mass and charge
state. Jets that are initially collimated flows of particles are likely to be splayed
into sheets of grains. When Cassini crosses one of these highly modified jets, it
may only encounter particles with a very limited range of M/q at a given location,
as may have occurred during encounter E10.
One puzzle arising from the CAPS observations is the colocation of positively
and negatively charged nanometer particles. The overriding charging process near
Enceladus is caused by the plume material’s immersion in Saturn’s corotational
magnetospheric plasma, resulting in negative potentials (Kempf et al 2006). The
co-existence of both populations at first appears anomalous. At nanometre scales,
however, the charging process can be stochastic, depending on local plasma parameters. In addition, during their subsurface formation in a collisional environment
(Schmidt et al 2008), grains could undergo triboelectric charging: those that condense within the vent even when of the same composition, can acquire opposite
charges. Smaller particles tend to charge negative, and larger ones positive (e.g.
Duff and Lacks (2008)). Most entrained nanodust particles are likely to have been
accelerated to near-gas velocities.
Overall, during both E3 and E5 encounters, CAPS detected negative particles to much lower kinetic energies than for positive particles; if a proxy for lower
masses, this supports the picture of triboelectric charging occurring within vents.
Although sunlight will have little effect in the near-Enceladus environment, particles’ charge state changes could vary once exposed to the plume and the corotational plasma flow, where plasma parameters can differ significantly. This may have
implications for the behaviour of particles under the influence of electromagnetic
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Sascha Kempf et al.
2062
forces, as their charge state changes, possibly reversing polarity, grains’ paths can
be affected significantly. Grains initially charged positive and deflected away from
Saturn may charge negative, causing them to then be accelerated in the opposite
direction. The positively-charged population that can be accelerated outwards by
the magnetospheric corotational electric field very likely forms a component of the
dust streams observed outside Saturn’s magnetosphere (Kempf et al 2005).
Farrell et al (2010) argued that the charge states observed by IMS and ELS
may not reflect the initial triboelectric charge state of the detected particles, due
to the likelihood of charge modification between their emergence at the surface
and their detection at Cassini. Shafiq et al (2011) also address the issue of grain
charging of particles exposed to the plume environment, based on RPWS Langmuir
Probe data. The susceptibility of low mass charged grains to electromagnetic forces
means that the fine structures in their signatures is expected to differ significantly
from those observed by CDA and RPWS, which detect high mass grains.
The presence of the dusty plasma envelope itself is thought to have a significant
effect on Saturnian magnetospheric plasma flow. Saturn’s magnetospheric plasma
is observed to slow down and be deflected in the vicinity of Enceladus (Tokar et al
2006). This is at least partially due to the mass-loading of the plasma by fresh
pickup ions created in and around the plume, but the presence of the charged
grains may also make a significant contribution to this deceleration and deflection
of the flow (Farrell et al 2010). The charged grains’ presence may have further
implications: Simon et al (2011) suggest that the nature of the electrodynamic
coupling between Enceladus and Saturn (Pryor et al 2011) is significantly affected
by the grains’ presence. Rather than being an observational curiosity, it appears
that this population of particles spanning the mass range between heavy molecules
and micron-scale dust has important effects on several processes occurring in Saturn’s inner magnetosphere.
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5.3 Electrodynamic coupling
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Satellites such as Enceladus that orbit within their parent planets magnetosphere
can interact electrodynamically with their surroundings, largely due to the relative
motion of the magnetospheric plasma and the moon. One of the key characteristics
of such interactions is a perturbation of the planetary magnetic field. Elements of
this interaction include the Pedersen and Hall currents generated by ion-neutral
collisions, and a pickup current generated by the incorporation of freshly-created
ions by the magnetospheric flow. These current are driven by the electric field that
drives the plasma E × B drift, and are perpendicular to the magnetic field. Closure
of these currents is achieved by currents flowing along the magnetic field, which
couples the interaction region in the vicinity of the moon to the ionosphere of the
planet. The moon-planet current system can be modelled as a standing Alfvén
wave propagating along the magnetic field to the planets ionosphere (Neubauer
1980). The locations at which these currents reach the planets ionosphere can be
observed as bright auroral footprints at Jupiter from radio to UV wavelengths,
where three of the Galilean satellites have a strong enough interaction for the precipitating currents to be observed remotely (Clarke et al 2002). Enceladus, with
its substantial plume and associated ionosphere, clearly has the ingredients for
a moon-magnetosphere interaction potentially strong enough to form remotelyobservable auroral spots at its two magnetic footpoints which may provide significant local energy input at the top of Saturns atmosphere. A search for such a
signature in HST UV images obtained between February 2005 and January 2007
was unsuccessful (Wannawichian et al 2008).
However, Pryor et al (2011) reported the detection of an auroral spot associated
with Enceladus in Saturns northern hemisphere. The search for this feature in
UVIS data was spurred on by the first detection of field-aligned electrons near
Enceladus, during the E4 encounter, when CAPS-ELS observed energetic (1 ∼
keV) field-aligned electrons arriving at Enceladus from the direction of Saturns
northern hemisphere, together with coincident energetic ions (55-90 keV) being
Enceladus as an active body
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observed in the same direction by MIMI. The beams were observed downstream of
Enceladus, from > 23.3 to 3.6 RE , where they stopped abruptly despite the viewing
geometry still being conducive to field-aligned observations. The beam electrons
flickered in energy between populations peaking at ∼10 eV and 1 keV; bimodal
populations were also briefly observed. These changes appear to be associated
with changes in the magnetic field, and hence represent a true change in the
field-aligned current density. Pryor et al (2011) estimated that the Alfvén wave
travel time between Enceladus and Saturns ionosphere would be 150 s, yielding a
distance for the beams to begin of ∼3.6 RE , consistent with the observations.
The auroral signature observed by UVIS was at 64.1 ± 0.4 degrees, near to
the predicted 64.5 degrees. The northern signature was not, however, always detectable, suggesting a significant variation in brightness, and its southern hemisphere counterpart was not observed at all. When observable, the northern spot
varied in brightness by a factor of ∼3. The spots size indicated an interaction
region extending to ∼20 RE downstream of the moon, and up to 20 RE in radial
width. The power transferred in this interaction is ∼ 3 · 108 W, several orders of
magnitude lower than the equivalent interactions at Io and Europa, at ∼1 012 and
∼1 011 W, respectively.
Further in situ observations of elements of the electrodynamic interaction
region were presented by Gurnett et al (2011), who reported the detection of
Whistler-mode auroral hiss emissions generated by magnetic field-aligned electron
beams from a few tens of eV to several keV in energy. Auroral hiss has a characteristic V-shaped feature when observed in frequency-time spectra obtained during
moon encounters, and was present in data from 6 of the 12 flybys studied. The
basic electrodynamic interaction at Enceladus is expected to consist of planet to
moon currents flowing on the Saturn-facing sides of the flux tube, and moon to
planet currents running on the anti-Saturn side. Saur et al (2007) noted a likely
important complication of this scenario at Enceladus, as the plume ionosphere to
the south of the moon is largely blocked from a direct connection to the northern
hemisphere due to the presence of Enceladus itself. They predicted the presence of
additional surface currents running between Saturns northern and southern hemispheres on the planet-facing side and in the reverse direction on the opposite side
of the interaction region.
Simon et al (2011) reported that the presence of charged dust in the vicinity
of Enceladus could significantly change the nature of the interaction, including
a reversal in the sign of the Hall conductivity. The same authors also reported
the detection in magnetometer data of the abovementioned blockage of the Alfvén
wing by Enceladus itself.
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5.3.1 Observations, models and particular aspects of Enceladus plasma interaction
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The plasma just upstream of Enceladus, but not in the plume itself (e.g., a radial
distance from Eneladus of some tens of RE ), as seen during the E2 flyby (Tokar
et al 2006) was observed by the CAPS IMS to consist mainly water group ions
(mostly H2 O+ , OH+ , and O+ with some H3 O+ . The measured drift speed is
what one would expect for corotation (ud ∼ 26 km s−1 in the Enceladus/Keplerian
frame of reference). The observed distribution function was torus-like (or ring-like
in velocity space) but with a rather large spread of velocities in the ring (Tokar
et al 2006). Ions created by ionization of neutrals, or by charge exchange reactions
with neutrals, are picked up by the magnetic and motional electric fields and
form a ring distribution function. The total observed plasma density upstream of
Enceladus is about 100 cm-3 (Persoon et al (2009a); Tokar et al (2006)).
Water group ions were observed in the plume itself by the CAPS IMS, and
the INMS experiments (Tokar et al (2006, 2009); Cravens et al (2009)). The total ion density appears to increase in the plume itself and the flow speed of the
ions relative to Enceladus was seen to becomes quite small (Tokar et al 2009).
Figure 19 shows ion count rate versus energy measured by CAPS IMS during the
E3 encounter. The ions move slowly with respect to Enceladus and are cold. An
approximate ion composition was obtained from CAPS data in this stagnation re-
52
Sascha Kempf et al.
Fig. 19 Cassini CAPS IMS count rate versus energy for ions during the close encounters of
E3 (top) and E5 (bottom).When the plasma is cold, the energy spectrum is approximately a
mass spectrum.Water group ions are evident in the plume. From (Tokar et al 2009).
2149
2150
2151
2152
2153
2154
2155
gion, again showing the presence of water group ions but with an enhanced H3 O+
abundance. The INMS in its ion mode observed ions that were almost at rest with
respect to Enceladus and the mass spectrum obtained showed right in the plume
(at a radial distance of a couple of RE ) that H3 O+ was at least ten times more
abundant than the other water group ion species (Figure 20). This species can be
explained by the chemical reaction (Young et al (2005); Tokar et al (2009); Delamere, P. A. and Bagenal, F. (2008); Cravens et al (2009); Sittler et al (2008))
2156
2157
H2 O+ + H2 O → H3 O+ + OH
2158
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H3 O+ has also observed to be the major species in the inner coma of active
comets due to this reaction (Balsiger et al 1986). The CAPS IMS and the INMS
also detected water cluster ions including H3 O+ - H2O and H2 O+ - H2 O (Cravens
et al 2009). CAPS ELS measurements have shown that negatively charged, as well
as positively charged, ions are present in the plume (Coates et al 2010b). The
range of ion masses is very large and extends up to nanometer sizes (i.e., grains)
(Jones et al 2009). These negative ions are poorly understood. The CAPS ELS
also measured the electron flux versus energy in the plume, as shown in Figure 21.
The electron distribution is dominated by electrons with energies of about 1020 eV. Ozak et al (2012) modeled the electron energy distribution in the plume
using a two-stream transport code that included the photoelectrons produced by
photoionization of neutrals by solar radiation, and elastic and inelastic (i.e., ionization collisions and electronic, vibrational and rotational excitation collisions).
The model-data comparsions (Figure 21) showed that most the electrons in the
plume are photoelectrons that are either locally produced or come from Saturns
magnetosphere outside the plume. The ledge in the CAPS ELS spectrum near
25 eV is due to photoionization of water by solar HeII resonant line photons and
such peaks are characteristic of photoelectrons in other space environments (see
Coates paper on Titan.).
Enceladus as an active body
53
Fig. 20 The ion composition measured by the Cassini Ion and Neutral Mass Spectrometer in
the plume during the E3 encounter. Mass 19 ions (H3O+) strongly dominate the composition.
From (Cravens et al 2009).
Fig. 21 The electron energy flux versus energy in the plume during the E3 encounter by the
CAPS ELS (blackline with dots). The spectrum modeled using a two-stream transport method
is shown as the yellow and red lines (no external electron inputs) or as the green and blue
lines (external electron inputs). The purple dashed line shows a Maxwellian distribution. From
(Ozak et al 2012).
2181
The external plasma flow has been observed to slow down as it enters the
plume, as discussed above. This is expected due to the mass loading of the flow
by ionization and charge exchange, which create pick-up ions. Numerical models
have been used to explain and interpret this interaction.
2182
5.4 E ring
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2184
There is an intimate connection between Enceladus and Saturn’s diffuse E ring,
which envelops the ice moons Mimas (3.07 RS ), Enceladus (3.95 RS ), Tethys (4.88 RS ),
54
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Sascha Kempf et al.
Dione (6.25 RS ), Rhea (8.73 RS ), and Titan ( RS ). The ring is composed of predominantly water ice grains (Hillier et al 2007) smaller than 10 µm (Nicholson
et al 1996; Kempf et al 2008). Numerical studies of the ring particle dynamics
suggest particle lifetimes of less than 200 years (Jurac 2001; Juhász et al 2007;
Beckmann 2008), implying that a mechanism resupplying the ring with fresh dust
must exist. Fourteen years after the discovery of the ring by ground-based observations (Feibelman 1967a), Enceladus was proposed to be the dominant source of
ring particles because the ring’s edge-on brightness profile has its maximum near
the moon’s orbital distance (Baum et al 1981). In infrared edge-on images taken
during Saturn’s ring plane crossing (August 1995), de Pater et al (2004) found
small density enhancements around the orbits of Tethys and Dione, suggesting
that all ring moons may contribute to the ring particle population. Surprisingly,
the de Pater et al (2004) finding has not been been confirmed by any of the Cassini
instruments (Horányi et al 2009). Only in situ data obtained by RPWS and CDA
during a close flyby of Rhea may provide evidence that a moon other than Enceladus injects a small amount of fresh dust into the ring (Jones et al 2008). The
current view is that Enceladus is by far the strongest, and probably the only,
source of E ring particles.
Two dust production mechanisms have been postulated for Enceladus. Interestingly, geyser-like processes were proposed early (Haff et al 1983; Pang et al 1984),
although in later years most attention was paid to meteoroidal impact ejection (e.g.
Hamilton and Burns (1997) or Juhász and Horányi (2002)). This change of focus
was probably due to the fact that the structure of Jupiter’s diffuse gossamer rings
could be explained by collisional ejecta originating from the ring moons Almathea
and Thebe, whose orbital extermes coincide with the outer boundaries of the rings
(Burns et al 1999). However, the discovery of the Enceladian dust jets dramatically
changed the picture, with a lot of the missing puzzle pieces falling into place. In
particular, two observations have intrigued scientists from early on. Firstly, the E
ring has a unusual blue colour, implying a narrow size distribution centred around
1 µm (Nicholson et al 1996), while rings resupplied with collisional ejecta would
be expected to have a broad size distribution. Secondly, the ring’s vertical scale
height is not bound by the orbital extremes of Enceladus but increases with the
radial distance to the moon (Showalter et al 1991b). This implies that either the
grains’ ejection speed is much larger than the satellite’s escape speed of 207 m s−1
, contradicting laboratory data for impact ejecta (Stöffler et al 1975; Hartmann
1985), or the inclinations of the ring particles have to grow rapidly after their
escape into the ring. Although the latter may be induced by the time-dependent
out-of-plane component of the solar radiation pressure (Horányi et al 1992), numerical simulations of the ring dynamics do not reproduce the full extent of the
observed vertical scale height (e.g. Juhász and Horányi 2002).
The injection of dust grains from localised sources in the south polar terrain
of Enceladus explains many aspects of the ring’s peculiar vertical profile. Since
the grains’ ejection speed is much slower than the Enceladus orbital speed vE ,
the orbital elements of fresh ring particles and the moon are almost identical.
However, because plume particles are ejected roughly orthogonally to the ring
plane, the minimum inclination of the ejected grains is imin = RHill / h rE i =
0.23◦ , where RHill = 948 km is the Hill radius of Enceladus, characterising the
range of moon’s gravitational influence, and h rE i = 3.95 RS is the moon’s mean
orbital distance from Saturn. Particles ejected faster than the three-body escape
speed from Enceladus of vesc = 207 m s−1 may escape to the ring and have an
initial orbit inclination of
1/2
2
v 2 − vesc
i ≈ 0.23◦ + e
.
(6)
vE
2236
2237
2238
2239
2240
Numerical studies by Kempf et al (2010) showed that, in order to escape to the
ring, plume particles need be launched faster than 224 m s−1 , as slower grains
recollide with Enceladus during their first orbit. Furthermore, the initial inclination
of grains launched slower than 244 m s−1 will be altered by three-body effects,
because these grains traverse the Enceladus Hill sphere during their first ring
Enceladus as an active body
55
0.10
Rate (1/s)
0.08
0.06
0.04
0.02
0.00
!0.10
!0.05
0.00
Height above ringplane (RS)
0.05
0.10
Fig. 22 Vertical profile of the E ring at Enceladus’ orbital distance as measured by the Cassini
dust detector CDA during a steep ring plane crossing at 3.93 RS (diamonds). The detector was
sensitive to grain sizes greater than 0.9 µm. The solid line compares the Cassini data with
model calculations for a ring fed by the 8 Enceladus dust jets identified by Spitale and Porco
(2007). Each pronounced spike in the vertical profile is associated with a dust jet emerging
from the moon’s south polar terrain (Kempf et al 2010).
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2273
plane crossings. By tracing the trajectories of plume particles numerically, Kempf
et al (2010) computed the spatial distribution of the plume particles in the vicinity
of Enceladus. The model includes the effects of the higher moments of Saturn’s
gravitational field as well the Lorentz force due to Saturn’s rotating magnetic
field acting on the charged grains. The time-dependent grain charge due to the
ambient plasma and the solar UV radiation is computed simultaneously with the
particles’ equations of motion. Finally, the distributions of the initial speeds and
grain sizes were as described by the Schmidt et al (2008) model. The resulting
vertical ring profile matches the measurements by the Cassini dust detector CDA
well. Figure 22 compares the simulations with CDA data obtained during a steep
Cassini traversal through the ring plane at approximately the orbital distance of
Enceladus (Kempf et al 2008). The kinks in the vertical profile are associated with
individual Enceladian dust jets and are reproduced well by the simulations.
The vertical ring profile is not the only imprint of the localised dust injection
by Enceladus’ south pole plume. The radial brightness profiles derived from earthbound observations indicated that the densest point of the E ring lies outside the
orbit of Enceladus, displaced by 104 km (de Pater et al 2004), while in situ measurements by Cassini CDA found the displacement to be > 3 000 km (Kempf et al
2008). Numerical simulations before the discovery of the Enceladian plume failed
to reproduce this observation. Juhász et al (2007) showed that the initial inclinations of the plume particles decrease the likelihood of the particles re-colliding
with Enceladus, allowing them to migrate outward by drag forces exerted by the
ambient plasma. The imprint of the Enceladian dust jets is also clearly visible in
Cassini images of the E ring (Fig. 23).
Despite the remarkable progress since the discovery of the Enceladus plume
there are many phenomena still not understood. One important open question is
whether the Enceladus geysers are really the only source of E ring dust. In situ
data imply a dust escape rate to the ring of 0.2 kg s−1 (Spahn et al 2006b), which
is surely enough to maintain a steady-state E ring. However, it is difficult to understand why the influx of interplanetary meteoroids does not produce noticeable
amounts of collisional ejecta. Spahn et al (2006a) reevaluated the expected ejecta
production rates of the E ring moons in the light of the new Cassini data. The authors conclude that impactors of interplanetary origin produce ∼ 0.1 kg s−1 , while
56
Sascha Kempf et al.
Fig. 23 Cassini image of the densest region of the E ring taken at a distance of about 2.1 ·
106 km from Enceladus (NASA/JPL/SCI PIA 08321). The moon is within a dust torus fed by
the Enceladian ice jets. The strands emerging from the moon form the various micro-signatures
apparent in the ring’s vertical (see e.g. Fig. 22) and radial profiles recorded by the Cassini in
situ instruments such as RPWS (Kurth et al 2006) or CDA (Kempf et al 2008). Tethys is
visible to the left of Enceladus.
2296
E ring impactors produce ∼ 10 kg s−1 of fresh dust. If this were the case, collisional
ejecta would still be the dominating E ring dust source. However, there is only
circumstantial evidence for ejecta in the Cassini data. The best fit of the CDA
plume density profile of grains ≥ 1.6 µm obtained during the first close Enceladus
flyby in 2005 (Spahn et al 2006b) to numerical plume models is obtained by assuming an additional ejecta source with a peak rate of 0.5 s−1 at closest approach
to Enceladus. A similar conclusion was reached by Spahn et al (2006b), who compared the CDA data with a plume formed from grains emerging from a circular
area of 30◦ centred at the moon’s south pole. If there is really an ejecta source
of this strength, then most of the ejecta grains must be produced by ring particle
collisions. In this case, however, ejecta from the ring moon Tethys should also form
a dust torus, which hasn’t yet been observed.
The large number of Cassini traversals through the inner E ring allowed the
Cassini in situ instruments, mainly RPWS and CDA, to monitor the time variation
of the ring particle density. Kurth et al (2006) reported on one equatorial and 5
vertical passages through the inner E ring during which RPWS measured ring
particle impacts, while Kempf et al (2008) published CDA dust density data for
grains ≥ 0.9 µm recorded during 2 equatorial and 6 vertical Cassini ring traversals.
The data by the latter team provides evidence for a variation in the ring’s peak
density of 20% within 37 days. It is not clear whether the observed variations
stem from azimuthal variations within the dust torus or from a time-variable dust
plume. Nevertheless, the in situ ring data indicate that the Enceladus dust source
supplies the ring with fresh dust at a surprisingly constant rate.
2297
5.5 Inner magnetosphere
2298
5.5.1 Energetic particle interaction
2274
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2302
Enceladus orbits Saturn within the planets radiation belts (Roussos et al 2011,
and references therein). The moons surface and plume material are therefore continuously bombarded with energetic particles. The interpretation of the energetic
particle signatures is complicated by the varying drift rates of charged particles
Enceladus as an active body
2303
2304
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57
in Saturns magnetosphere. All protons and other positive ions, plus low energy
electrons, have a net drift motion around Saturn in the same sense as the planets rotation. In the rest frame of the corotating thermal plasma, electrons drift
opposite to the direction of planetary rotation. At energies around 1 MeV, this
counterdrifting motion is of a large enough magnitude for electrons to reverse direction in the inertial frame too: electrons of > 1 MeV travel around Saturn in
the direction opposite to the planets rotation. In the Enceladus frame, the lowenergy thermal plasma is continuously overtaking the moon: the moons trailing
hemisphere is therefore exposed to bombardment by low energy ions and electrons. The bulk motion of energetic electrons is however in the opposite sense, so
Enceladuss leading hemisphere is most exposed to these particles.
As is seen in the vicinity of the orbits of other large icy moons, the absorption
of energetic particles leads to the formation of narrow cavities of reduced energetic
particle number densities. Due to the rapid bounce motion of energetic particles
(>tens of keV), these cavities extend to higher latitudes, well out of the plane of the
moons orbit. On crossing these cavities, instruments on Cassini detect dropouts
in energetic particle fluxes; these are seen as a drop in foreground particles by
MIMI-LEMMS, but also cause background decreases in certain CAPS and UVIS
data. Such cavities that persist over limited longitude ranges and time periods are
referred to as microsignatures. First observed by Pioneer 11 and the Voyagers at
the L shells of several icy moons (van Allen et al 1980; Carbary et al 1983), these
signatures were successfully modelled in detail by Paranicas and Cheng (1997).
On arrival at the Saturn system, such signatures were clearly observed by MIMILEMMS on numerous occasions (Paranicas et al 2005; Jones et al 2006).
Due to magnetospheric particle diffusion, the microsignatures gradually refill,
and those of Enceladus are usually eroded while being carried around Saturn before
returning to the location of the moon. The energetic electron microsignatures that
form upstream of Enceladus, at energies greater than ∼ 1 MeV, are seen to be
relatively sharp and, as they match the diameter of Enceladus in scale, are largely
unaffected by the presence of the plume (Jones et al 2006).
Ahead of the moon in the Keplerian sense, or downstream of the moon in terms
of the corotation flow, is a wake of low-energy ions and electron. To date, Cassini
has passed through this region during encounters E3 to E6 inclusive. In CAPSELS data, a clear absence of thermal electrons exists during the wake traversal. At
slightly higher energies, i.e. tens of keV or more, where electrons execute significant
bounce motion, the wake extends north and south from the moons orbital plane.
The profiles of the microsignatures at there energies are generally less sharplydefined than the high energy microsignatures seen upstream of the moon, and
exhibit widths in excess of the diameter of Enceladus. It is believed that these
intermediate-energy electrons are partially absorbed by the plume. Variability in
the depth and width of these signatures is interpreted as evidence for variability
in the plume density, and hence in its activity level (Jones et al 2006).
As energetic ions drift around the planet at such a high rate, the rate of absorption by the moon can be comparable to the rate at which new ions are created
or drift into the absorption region. Persistent depletions termed macrosignatures
are observed in ions of several MeV (Jones et al 2006). The rate at which microsignatures refill, and shift in radial position, can reveal significant information
on the magnetosphere (e.g. Roussos et al (2007)).
The exposure of Enceladuss surface to the energetic particle environment has
observable effects. Early works, e.g. Lanzerotti et al (1983), considered the production of a neutral torus encircling Saturn that consists of material sputtered from
the surfaces of Enceladus and other icy moons. The sputtering rates resulting from
this exposure to energetic particles have been calculated in detail (Johnson et al
2008); it should be noted that this process affects plume and E-ring grains as well
as the moons surface ices. The energy spectra that the surface of Enceladus is
exposed to was computed by Paranicas et al. (in press), based on observations
made by the MIMI instrument. Although not yet studied in detail for Enceladus,
we note the possible importance of surface charging (Roussos et al 2010); at Enceladuss location in the inner magnetosphere, incident plasma will generally tend to
58
Sascha Kempf et al.
2376
charge the moons surface negative, overwhelming photoemission that charges the
surface to a positive potential.
Enceladus exhibits a subtle albedo asymmetry (Buratti et al 1990; Verbiscer
and Veverka 1994; Schenk et al 2011), and generally displays stronger water absorption bands on the Keplerian-leading hemisphere (Verbiscer et al 2006), at least
partially due to the irradiation of its surface. Clobal colour mapping also reveals
the presence a significant colour asymmetry in the form of depressed IR/UV ratios, i.e. more bluish terrain, centered on the south pole and extending northwards
at two longitude bands at ∼40 W and ∼220 W. The colour patterns are distinct
from neighbouring large icy satellites; no darkening of the trailing hemisphere is
observed. It is likely that the possible near-continuous deposition of plume grains
on the surface is responsible for this pattern (Kempf et al 2010), largely masking the effects of irradiation. The irradiation of the surface by magnetospheric
energetic particles could transform crystalline surface ice to an amorphous state
(Brown et al 2006).
2377
5.5.2 Neutral and plasma tori
2362
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2406
The emissions of Enceladus create the E-ring and a cloud of neutrals that form a
torus centered on the orbit of Enceladus. This neutral cloud then supplies most of
the ions in Saturns inner magnetosphere. In this section we discuss the neutrals
and low-energy ions in the inner magnetosphere and provide an overview of their
transport and loss processes.
The first strong evidence of a neutral cloud was a neutral-water dissociation
product, OH, observed from 3 to 8 RS in Hubble Space Telescope (HST) spectral
observations. Shemansky et al (1993) reported these hydroxyl neutral molecules,
which indicated the presence of water, in quantities more than predicted by theoretical estimates of sputtering of the rings and icy satellites (Ip 1997; Jurac et al
2002). Later observations and modeling by Jurac (2001) showed that OH density peaked near 4 RS , suggesting that Enceladus could be a source of H2 O for
the Saturn system. The OH observations with HST were sparse, and building
sufficient statistics required summing data from several observation periods. The
initial models of OH and its parent, H2 O, identified a radial and vertical (normal to
the equatorial plane) structure to the neutrals, but assumed azimuthal symmetry
(Richardson et al 1998; Jurac 2001; Johnson et al 2006).
Evidence of IM ions predates the discovery of the neutral cloud. During the
Voyager missions, the Voyager Plasma Science experiment data showed ions with a
mass greater than H2 in Saturns inner magnetosphere (Lazarus and McNutt 1983).
The energy-per-charge measurements were insufficient for unambiguous identification, but the nitrogen atmosphere at Titan suggested that these heavy ions were
N+ . The neutral OH by Shemansky et al (1993) indicated that the heavy ions were
likely to be water-group ions. Cassini instruments have confirmed this identification with remote UV measurements of neutral O molecules (Esposito et al 2005),
in situ ion measurements of water-group ions (Tokar et al 2006; Young et al 2005),
the discovery of water vapor emanating in plumes from Enceladus (Dougherty et al
2006; Hansen et al 2006; Porco et al 2006; Waite et al 2006), and the measurement
of H3 O+ ions in the plume (Cravens et al 2009).
2407
Transport and loss When neutral molecules leave the surface of Enceladus, they
2378
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2416
2417
become individual particles in orbit around Saturn. For some of the larger jets,
the gas may be collisional for 10s of meters, but then becomes collisionless. There
are several possible fates of these particles:
1. Collision with Enceladus or some other icy satellite. Although it is likely
that vapor leaves the surface of Enceladus with different velocities (see Chapter 2), most of the vapor will have speeds greater than the escape velocity, about
240 m/s. (Lower velocities will fall directly back to the surface.) However, as
described in the E-ring discussion, molecules that initially escape Enceladus
may impact Enceladus within a few orbits, usually within an integer multiple
of 16.5 hours, half of Enceladus orbital period. After dispersing away from
Enceladus as an active body
2418
2419
2420
2421
2422
2423
2424
2425
2426
2427
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2429
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59
Enceladus, particles may be swept up by other moons or by the rings. Indeed,
at 180 000 km, the inner boundary of the neutral cloud coincides with the G
ring, which is the outermost ring (excluding dust rings).
2. Dissociation. Electrons, ions, photons, and, less frequently, other neutrals,
dissociate neutral molecules into smaller molecules or atoms. Sometimes, ionization also causes dissociation, creating a neutral product along with the ion.
This fate changes the composition of the neutral cloud but does not reduce its
density.
3. Charge exchange with an ion in the magnetosphere
4. Ionization (excluding charge exchange)
5. Ejected from the Saturn system
Several processes contribute to dispersing neutral molecules. Many are associated with the processes already discussed at loss mechanisms.
1. Initial velocity. A molecule with a velocity of 500 m/s above escape velocity
will reach 10,000 km above and below the equatorial plane. Molecules with just
100 m/s excess velocity in the along-track direction will drift with respect to
Enceladus.
2. Neutral-neutral collisions. Originally considered negligible, Farmer (2009)
showed the importance of this process and subsequent modeling confirmed
that neutral-neutral collisions contribute at least half of the dispersion of the
neutral cloud. The mean free path is large, 104 to 105 km, but there are still
several collisions each orbit around Saturn. These collusions can contribute
nearly all of the
3. Velocity associated with ionization byproduct.
4. Charge exchange
2445
Cassini observations in 2005 (Dougherty et al 2006; Hansen et al 2006; Waite et al
2006) discovered water emanating from the south pole of Enceladus, verifying
Enceladus as the source of the observed OH and predicted H2 O neutrals.
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Distribution The distribution of ions and neutrals in Saturns IM and their dis-
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tribution within the water group is sensitive to loss and transport assumptions,
but there are direct measurements that provide calibration points for the models.
Supporting direct observations of O, OH, and H2O remote from Enceladus include
CAPS data (ions) averaged over several orbits and a few INMS observations of
neutrals and ions. There are observations limited to a single species, including
UVIS measurements of O and the HST measurements of OH (Shemansky et al
1993). The resolution of measurements from CAPS is barely sufficient to quantify
the separate ion species within the water group. Modeling (Jurac 2001; Smith et al
2010) implies a distribution amongst water-group neutrals. Additional transport
mechanisms affect the ions, so their distribution is less well determined by models.
Since source of the neutrals is highly localized, the neutrals are unlikely to
be azimuthally symmetric. Recent models are investigating the geometric dependence of the neutrals (Leisner 2009; Smith et al 2010), but the data available are
insufficient to constrain the azimuthal distribution in the models.
The ion composition of the plasma in either the torus or plume can be altered
from its primary state (i.e., H2 O+ , OH+ ) by ion-neutral collisions (i.e., charge
exchange and chemical reactions). Prior to the Cassini mission models of the inner
magnetosphere (e.g., Scarf, F. L. and Frank, L. A. and Gurnett, D. A. and Lanzerotti, L. J. and Lazarus, A. and Sittler, Jr., E. C. (1984); Lazarus and McNutt
(1983); Richardson et al (1998) predicted the dominance of water group ions for
the ion composition. More recently, Fleshman et al (2010) and Delamere, P. A.
and Bagenal, F. (2008) have modeled the composition and energetics of the Enceladus torus, pointing out the importance of ion-neutral collisions and chemistry.
The CAPS IMS experiments demonstrated that water group ions dominate the
composition with the most abundant species being H2 O+ and O+ (Jr. et al (2006);
Wilson et al (2008); Young et al (2005); Thomsen, M. F., Reisenfeld, D. B. ,Delapp, D. M., Tokar, R. L., Young, D. T., Crary, F. J., Sittler, E. C., McGraw,
60
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Sascha Kempf et al.
M. A. Williams, J. D. (2010)). H3 O+ ions are present in the torus, but at relatively lower densities compared with the plume plasma (Wilson et al (2008); Tokar
et al (2006)).
The ion distribution function observed in the Enceladus torus appear to be
pancake-like in velocity space (Wilson et al (2008); Tokar et al (2006)). The ions
drift at roughly the co-rotation speed (∼45 km s−1 ), e.g., Figure 24. The perpendicular average thermal velocity is about 20-30 km s−1 with the parallel velocities
being less than perpendicular velocities. Ions are initially created at rest with respect the neutrals (moving at close to the Keplerian speed) but then they are
picked up by the magnetospheric motional electric field and the magnetic field
forming a ring distribution. But given that a simple ring distribution in velocity space is not observed, the ion distribution functions must evolve (or partially
thermalize) due to Coulomb collisions and/or wave-particle interactions. Note that
charge exchange collisions remove ions but replace them with ions that are at rest
in the neutral frame of reference. This has the effect of removing momentum from
the ion population and increasing the ion thermal energy. The electron energy
distribution measured in the Enceladus torus, like that in the plume, appears to
be predominantly consist of few eV electrons (Sittler, E. C., Ogilvie, K. W., Scudder, J. D. (1983); Young et al (2005); Rymer, A. M., Mauk, B. H., Hill, T. W.,
Paranicas, C., Andre, N., Sittler, E. C., Mitchell, D. G., Smith, H. T., Johnson,
R. E., Coates, A. J., Young, D. T., Bolton, S. J., Thomsen, M. F., Dougherty,
M. K. (2007); Schippers, P., Andre, N., Johnson, R. E., Blanc, M., (Dandouras),
I., Coates, A. J., Krimigis, S. M., Young, D. T. (2009); Delamere, P. A. and Bagenal, F. (2008)). (Schippers, P., Andre, N., Johnson, R. E., Blanc, M., (Dandouras),
I., Coates, A. J., Krimigis, S. M., Young, D. T. 2009) fit a typical CAPS ELS electron spectrum (at r∼5 RS ). ) with three populations:
1. a cold approximately Maxwellian core
2. a photoelectron component (explaining the observed ledge near an energy of
25 eV- as discussed earlier, such a feature is also present in the plume electron
spectrum)
3. a more energetic tail with energies in excess of 50-100 eV.
2518
The fluxes for the energetic tail were quite small however. RPWS Langmuir
probe data demonstrated that the core population has a temperature of Te ∼13 eV near 4 RS (Gustafson, G., Wahlund, J. E. 2010). Cravens, T. E., Ozak, M. S.,
Campbell, I. P., Perry, M., Rymer A. M. (2011) numerically modeled both the
suprathermal electrons and the colder core electrons. The suprathermal electron
flux in the torus appears to be very similar to that in the plume (i.e., Figure 21)
.The heating of thermal electrons is mostly due Coulomb collisions with hot pick-up
ions (Sittler, E. C., Ogilvie, K. W., Scudder, J. D. 1983; Gustafson, G., Wahlund,
J. E. 2010) but Coulomb collisions with photoelectrons contributes about 20%
(Cravens, T. E., Ozak, M. S., Campbell, I. P., Perry, M., Rymer A. M. 2011). The
core electrons are cooled mainly by inelastic collisions with water (i.e., rotational
and vibrational excitation) (Gustafson, G., Wahlund, J. E. 2010; Cravens, T. E.,
Ozak, M. S., Campbell, I. P., Perry, M., Rymer A. M. 2011; Cravens, T. E., Ozak,
A. 1986).
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5.5.3 Low energy electron enhancements
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Low energy electron features, termed spikes, have been observed as discrete, short
duration enhancements in the low energy plasma density. Their locations suggest
an association with Enceladuss L shell and have been observed over a wide range of
latitudes. The spikes reach over 100 eV in energy and length scales up to hundreds
of kilometres. They have been observed continually since 2005 so do not appear
to be seasonal. There are no obvious enhancements parallel to the magnetic field
direction, signifying that that the electrons responsible for the spikes are not necessarily field aligned. The spikes appear to be more prevalent outside Enceladuss
L shell and tend to be more common near the moon in longitude. Some of these
spikes can be resolved into dispersed features when viewed in data that has had
Enceladus as an active body
61
Fig. 24 Ion distribution function measured by the CAPS IMS in the Enceladus torus in the
corotating frame of reference. Perpendicular thermal speeds exceed parallel thermal speeds.
From (Tokar et al 2008).
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the penetrating radiation removed. The spikes have varying characteristics; some
show energy dispersion, sometimes the peak energy differs, some are single blobs
whereas others are clusters of discrete spikes. This suggests that several formation
mechanisms may cause these signatures. The fact that spikes cluster around the
moons L shell strongly suggests that at least some of them are linked to Enceladus.
The process(es) by which these features are formed, and the cause for their
creation, are unidentified at present. However, there are several under consideration. Spacecraft potential changes, hence spacecraft charging, due to energetic flux
variations associated with microsignatures (Khurana et al 1987) have been ruled
out, as the spikes occur before and after microsignatures.
The surface of Enceladus is very likely charged; the surface charging of Saturnian icy moons has been address by Roussos et al (2010), and is expected to
have a negative potential due to the high density of the plasma impinging upon
it. The degree of negative charging varies across the surface, and changes as the
moon orbits Saturn due to the direction of incident sunlight changing (Roussos
et al 2010). Electrons flowing along magnetic field lines connected to the moon
encounter this negative potential and, if their energy is less than the magnitude of
62
Sascha Kempf et al.
2574
the negative potential, are reflected. Secondary electrons are also formed from the
acceleration due to the electric field near the moon’s surface. The electron beam
would have an energy corresponding to the moon’s surface potential and electrons
would only be reflected when their energies are less than the surface potential.
Surface charging has been detected at another kronian moon, Rhea (Jones et al.,
in preparation).
The low energy electron enhancements may result from electron reflection from
Enceladuss surface or even at the plume itself. Negatively charged plume grains
may collectively reflect incident low energy electrons. However the potential levels implied by the spikes peak energies do seem higher than the levels expected;
Enceladuss plume is expected to charge negatively only to about -2 V.
Spikes have also been seen in relation to injection events (DeJong et al 2010).
Another possible cause for the spikes may be related to the moon and plumes absorption of a portion of the local plasma population. This may indicate a plasma
instability that causes a small-scale injection event; i.e. the moon itself or the presence of its wake may be responsible for the formation of these dispersion features.
Either the dispersion features are created at the moon L shell or the plasma instability causes a flux tube to move inwards without the corresponding outward
movement of another flux tube. One limitation of looking for these dispersion
events at 4RS is that identification becomes challenging due to high energetic
electron background levels as we travel closer and closer to the planet Saturn.
Finally, some of the spikes could be fossil’ signatures of the moon-planet electrodynamic interaction. Supporting this is the fact that a few of the electron features
are seen to extend up to high energies, more than 3 keV, and are dispersed in
energy, indicating an age of several hours or more. The electrodynamic interaction
is known to accelerate electrons (Pryor et al 2011), but once the interaction ceases
to operate, the accelerated electron population may drift near Enceladuss L shell
where it remains observable for several hours or more.
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5.6 Outer magnetosphere and solar wind
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Enceladus plays a major role in controlling global magnetospheric dynamics at
Saturn. Contrary to some expectations prior to Cassinis arrival at Saturn, it appears that Titans role as a source of neutrals and ions in planets magnetosphere is
relatively small: few N+ ions are observed. The main ring system too is a negligible
source of plasma, as the population of O2 + expected from a ring source is weak.
Finally, the appears to be little input of material from the solar wind in the form
of alpha particles, H2 + and H3 + . This leaves Enceladus as the primary source
of magnetospheric plasma, both directly, from the ionization of plume material in
the moons immediate vicinity, and indirectly, through the ionization of the neutral
torus that results from Enceladuss expulsion of neutral gas and material liberated
from E ring grains. The neutral torus has its peak number density at the orbit of
Enceladus; these neutrals spread radially inwards and outwards, as evidenced by
the OH profile observed by HST (Johnson et al 2006). In situ CAPS observations
show that water ions are the dominant ion species out to 20 Saturn radii from the
planet.
The distribution of magnetospheric plasma is governed by equilibrium diffusion (Persoon et al 2009b). In the model of Persoon et al (2009b), which fits the
observations well, the effect of gravity is negligible, and the diffusion is dominated
for water group ions by the centrifugal force. The ambipolar electric field is significant for protons <7 RS from the planet, and magnetic mirroring is dominant for
the same hydrogen ions at greater than 7 RS . The model predicts that cold water
group ions are confined to low latitudes by the centrifugal force. The addition of
water group ions to the rotating magnetosphere leads to the deceleration of the
flow at the radial distance of Enceladus (Wilson et al 2009).
The mass-loading of the inner and middle magnetosphere leads to this region
being unstable to the radial interchange of magnetic flux tubes. Magnetic flux
tubes carrying cold, heavy ions migrate outwards to be replace by hotter, low
Enceladus as an active body
63
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density plasma moving towards Saturn (e.g. Hill et al (2005); Burch et al (2005));
this process has been modelled (Hill 2008).
2605
5.6.1 Oxygen at Titan
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We note that material originating at Enceladus may have significant effects on
the atmospheric chemistry occurring at Titan: Sittler et al. (2009) proposed that,
as described above, the dissociation and ionization of water vapour in the plume
leads to the formation and ionization of oxygen in the magnetosphere. The flux of
O+ at Titan is estimated to be ∼1·106 ion cm-2 s-1 (Hartle et al 2006). The vast
majority of this O+ is believed to go towards the formation of CO or CO2 + in
Titans atmosphere, but 0.1% can be incorporated into cage-like molecules such
as C60 fullerenes. Heavy aerosol particles fall through the atmosphere, forming
haze layers, and they eventually reach the surface, where the oxygen is available
to partake in chemical reactions. Sittler and colleagues suggest that the presence
of oxygen atoms on Titan, originating at Enceladus, could therefore provide the
basis for pre-biological chemistry.
64
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http://adsabs.harvard.edu/abs/1984satn.book..318S@INBOOK1984satn.book..318S, author = Scarf, F. L. and Frank, L. A. and Gurnett, D. A. and Lanzerotti, L. J.
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DIAGNOSTICS, SATURN ATMOSPHERE, SPACE PLASMAS, ELECTROSTATICS,
PIONEER 11 SPACE PROBE, PLASMA LAYERS, PLASMA WAVES, RADIATION BELTS, VOYAGER PROJECT, WAVE INTERACTION, booktitle = Saturn, year = 1984, editor = ”Gehrels, T. & Matthews, M. S.”, pages = 318-353,
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