Fissure eruptions in Tharsis, Mars: Implications for eruption conditions and magma sources

Journal of Volcanology and Geothermal Research 185 (2009) 28–46

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Journal of Volcanology and Geothermal Research

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Fissure eruptions in Tharsis, Mars: Implications for eruption conditionsand magma sources

Lionel Wilson a,b,⁎, Peter J. Mouginis-Mark a, Shelly Tyson b, Jennifer Mackown b, Harold Garbeil a

a Hawai'i Institute of Geophysics and Planetology, University of Hawai'i at Manoa, Honolulu, Hawai'i 96822, USAb Lancaster Environment Centre, Lancaster University, Lancaster LA1 4YQ, UK

⁎ Corresponding author. Hawai'i Institute of Geophysof Hawai'i at Manoa, Honolulu, Hawai'i 96822, USA.

E-mail address: [email protected] (L. Wilson)

0377-0273/$ – see front matter © 2009 Elsevier B.V. Adoi:10.1016/j.jvolgeores.2009.03.006

a b s t r a c t
a r t i c l e i n f o
Article history:Received 12 June 2008Accepted 10 March 2009Available online 26 March 2009

Keywords:fissure eruptionsTharsis, Marseruption ratesfire fountainsspatter ramparts

The Tharsis region has been the focus of many studies of volcanism on Mars. Not only are the largestvolcanoes on the planet located here, but also many of the youngest volcanic features are found withinTharsis. Comparatively little attention has been given to understanding the origin of the smaller volcanoeswithin Tharsis, and these smaller structures may provide important information on both the regionaltectonics of the area, and the availability of magma from the shallow or deep crust. We have identified sixseparate, but physically close, vent systems in eastern Tharsis just to the east of the volcano Jovis Tholus.These vents are typically linear fissures a few to ~20 km in length that have built small shields rising to ~50–85 m above the level of the surrounding topography. The maximum length of individual lava flows fromthese fissures is ~30 km, but more typical lengths are 15–20 km. Mapping of the vent complexes reveals thatsmooth material, interpreted to be spatter from fire fountaining, is located on the rims of some of the fissures.This spatter then formed either short (b5 km) flows or merged close to the fissure to form the longer flows.There appears to have been some temporal evolution of the flow field, as the oldest parts of the basement ateach center are built from a series of compound flows that cannot be subdivided into individual flows. Wefind evidence of frozen lava ponds and channelized flows from central vents. Photoclinometric profiles areconstructed across one of the longer flows to confirm an unusual attribute, originally identified by Mouginis-Mark and Christensen [Mouginis-Mark, P.J., Christensen, P.R., 2005. New observations of volcanic features onMars from the THEMIS instrument. J. Geophys. Res. 110: E8, doi: 10.1029/2005JE002421], of the flows fromthe central vent complex: they are all b5 m thick, which is a factor of ~8–15 thinner than flows elsewhere onMars. To investigate the unusual eruption conditions (limited total volume of the construct and of individualflows) we model the lava rheology, the duration of emplacement, the subsurface conditions that may haveled to these eruptions, and the volatile content of the magma. We find lava viscosities and yield strengths tobe ~100 Pa s and ~100 Pa, respectively, eruption rates to be ~5000 m3 s−1, flow speeds to be 1–2 m s−1,durations of emplacement of individual flow units to be ~5 h, and the equivalent magma water content to be0.1–0.2 mass%. These eruption conditions are consistent with a wide range of possible depths of the magmareservoirs feeding the eruptions. Small vents such as the ones studied here have also been identified in otherparts of eastern Tharsis, so that the eruptions described here may characterize a common style of volcanismon Mars that can only be identified now that image spatial resolution in the range 1–20 m/pixel is available.It is therefore possible that additional searches of other volcanic areas (Syrtis Planum, Elysium Planitia, andHesperia Planum) may also show greater diversity of activity than is currently accepted.

© 2009 Elsevier B.V. All rights reserved.

1. Introduction

Volcanic landforms and eruptive processes on Mars have beenstudied since the 1960s when the Mariner missions returned imagesof the surface showing extensive modification by volcanism (McCau-ley et al., 1972; Carr, 1974). Since that time, numerous investigations

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have been conducted on many aspects of Martian volcanism such asvolcanic features and landforms (e.g., Greeley and Spudis, 1981;Wilson and Head, 1994; Mouginis-Mark and Christensen, 2005)eruptive processes (e.g., Wilson and Head, 1994; Sakimoto et al.,1997), magma sources (e.g., Wilson et al., 2001; Wilson and Head,2002) and lava flow rheology (e.g., Baloga et al., 2003; Rowland et al.,2004; Glaze and Baloga, 2006). However, few of these investigationshave concentrated on small-scale volcanic processes due to the lack ofimage data at a sufficient spatial resolution.

It is generally accepted that levels of volcanic activity havedecreased with time onMars; as the lithosphere cooled and thickened

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Fig. 1. Location map of the study area within NE Tharsis. The volcano Olympus Mons isat top left of the image, Ascraeus Mons is at far right, and Pavonis Mons is at lower right.The position of the fissures studied here to the east of Jovis Tholus volcano is illustratedby the white rectangle, which delineates the coverage of Fig. 2. Base image is a shadedrelief image derived fromMOLA data, and covers the area from 2.5°S to 25.0°N, 220.0°Eto 259.0°E.

29L. Wilson et al. / Journal of Volcanology and Geothermal Research 185 (2009) 28–46

the style of volcanism evolved. Initially vast outpourings of lava camefrom linear fissure systems, flooding parts of the ancient crateredsurface to form the volcanic plains (Greeley and Spudis, 1981). As thelithosphere thickened it restricted the ascent of large volumes of melt,thereby reducing the supply of magma to the surface. The Tharsisvolcanic province is thought to have resulted from a long-livedmantleplume producing episodic batches of melt (Wilson et al., 2001). Thesediapirs rose through the crust to a shallow depth where they achievedneutral buoyancy (Wilson and Head, 1994). Eruptions at the surfacewere fed by dikes from these high-level “magma chambers”constructing central edifices such as the Tharsis Montes, as well asthe smaller volcanic constructs such as the tholi.

The smaller (b~50 km in diameter) volcanic features withinTharsis may provide valuable information on the distribution ofmagma in the shallow crust towards the late stages of volcanism in thearea, and on the evolution from point-source vents of activity to theconstruction of the tholi and montes. They may also providesupplemental information on the tectonic evolution of the Tharsisregion (e.g., Anderson et al., 2001). In addition, by virtue of theirrelatively small size, it may also be possible to place constraints onindividual eruptions because entire flow fields can be mapped. Since2001, the Thermal Emission Imaging System (THEMIS) aboard theMars Odyssey spacecraft has been collecting a large data base of 18-meter per pixel images with a wide spatial coverage (Christensenet al., 2004). THEMIS visible (VIS) images and data from the HighResolution Stereo Camera (HRSC) have enabled individual long lavaflows to be mapped (Baloga et al., 2003; Hiesinger et al., 2007), butuntil now the distribution of vent-specific flows has not been studied.The continuing Mars Odyssey mission has now collected a sufficientnumber of THEMIS VIS images to permit us to identify three ventcomplexes that collectively form part of a larger systemwhich extends~45–240 km to the east of the small volcano Jovis Tholus.

Fig. 2. (a, Upper image). Uncontrolled photomosaic of 46 individual THEMIS VIS images of thFigs. 5a, 6a, 6b, 6c, 7, 8, and 9 are outlined. (b, Lower image). Sketch map of the study area, baThe four areas of recent activity considered here (Figs. 3 and 4) are shaded, and the three sep

2. Description of vents

Our study area comprises a portion of NE Tharsis extending from2.5°S to 25.0°N, 220.0°E to 259.0°E (Fig. 1). Our mapping has identifiedthree vent complexes (Fig. 2a), labeled Western vents, Central ventsand Eastern vents (Fig. 2b), each of which comprises a fissure that hasbuilt a low shield and erupted lava flows that extend ~20–25 km fromthe fissure (Fig. 2b). Within the easternmost of these three vent

e study area, which extends from 17.3° to 19.0°N, and 241.5° to 247.5°E. The locations ofsed on the photomosaic shown in Fig. 2a. Jovis Tholus is the circular construct at far left.arate vents in the eastern complex are individually identified by the numbers 1, 2 and 3.

Table 1Geometries of eruption centers. Values are the estimated maximum heights, surfaceareas, and volumes for the five main eruption centers identified in Fig. 2 b. Maximumheight is defined as the maximum height of a shield above the average of thesurrounding terrain at the perimeter of the shield.

Shield name Height(m)

Area(km2)

Volume(km3)

West vent 180 2160 111Central vent 75 1110 27East vent #1 60 501 7East vent #2 43 124 1.3East vent #3 150 1440 39

30 L. Wilson et al. / Journal of Volcanology and Geothermal Research 185 (2009) 28–46

complexes there are three separate centers (Fig. 2b) denoted, fromwest to east, Eastern vent #1, Eastern vent #2 and Eastern vent #3. Anadditional area of recent lava flows can also be identified within alarge (~18-km diameter) impact crater located just to the east of theCentral vent complex at 18.3°N, 246.0°E; we regard this as part of theCentral vent system.

Using the 128th-degree MOLA digital elevation model (DEM), wehave calculated the surface area and volume of each of the five mainvents identified by this mapping. The surface area of each vent systemwas determined via conventional photogeologic mapping to establishthe perimeter. The volume was found by fitting a two-dimensionalsecond-order polynomial to the perimeter to approximate the pre-eruption surface, summing the height differences between the DEMsurface and the fitted surface, and multiplying this sum by the~463×463 m2 DEM pixel area. The volumes range from 1.3 km3 forthe lava pond that forms vent #2 in the eastern vent complex to111 km3 for the western vent complex (Table 1).

3. Styles of eruption

We have mapped the three largest vent complexes to identify thestyle(s) of eruption (Fig. 3). There are several similar features at thesethree locations, as well as some unique site-by-site attributes. Ingeneral terms, the common attributes of each vent system include:

Fig. 3. (a, left) Interpretative map of the lava flow associated with the Western vent complexorange and light brown. Dark green flows are prominent flows that originate from the zoneburied but probably also originated from the zone of spatter. The blue area marks the extentmargins can be identified. Contours, derived from the 128th-degreeMOLA digital elevation mat a spatial resolution of 18m/pixel. (b, center) Interpretativemap of the lava flowassociatedas Fig. 3a with the addition of the purple unit, which appears to be an earlier, partially buriefrom the 128th-degree MOLA digital elevation model, are at 25-m intervals relative to the Mright) Interpretative map of the lava flows associated with the Eastern vent complex. See Figboundaries between vents 1, 2 and 3 are denoted by the thick black lines. Color scheme of unimodel, are at 25-m intervals relative to the Mars datum. Mapped from THEMIS VIS images atfigure legend, the reader is referred to the web version of this article.)

(1) a 330–1100-m-wide linear fracture; (2) an area of smooth terrainon each side of the fracture; (3) short (b8 km), narrow (b2 km) flowsthat originate from this smooth terrain; (4) longer (10–25 km) flows,many of which originate from the fracturewhereas others are partiallyburied at their proximal ends by more recent flows; and (5) an area ofundifferentiated flows that form the perimeter of each vent complex.In a few areas, such as the eastern vent complex, we can identify small(b5 km wide) patches of an earlier phase of these undifferentiatedflows.

The Western vent complex (centered at 18.0°N, 243.6°E) is builtaround a 10.6-km-long fissure (Fig. 3a). This vent system appears tohave been most productive at its southern end. The Central ventcomplex (originally discussed by Mouginis-Mark and Christensen(2005)) is centered at 18.1°N, 245.0°E (Fig. 3b). It comprises a 43-km-long series of collapse craters, of which the central 17 km has mergedinto a single fissure ~350 m wide. At both ends of the vent complex,we find collapse pits that are aligned with the strike of the fissure butwere not associated with the eruption of lava. At the western end,there are 16 individual pits in a 20-km segment of the fissure system,with three isolated pits that are ~6.2 kmwest of any identified sourceof lava. At the eastern end of the fissure, the pits occur within theejecta blanket of an 18.7-km diameter impact crater. We interpretthese pits to be collapse features above a subsurface dike that did noterupt any magma at these locations.

Just to the east of the central vent complex, the fissure appears tohave become productive again, this time inside an impact crater(Fig. 4). No signs of a vent can be found within the crater, but the floorof the crater is very smooth, there are flow lobes (approximatelyconcentric with respect to the crater rim) on the crater floor, and thereis a low circumferential bench that could be formed by drain-back oflava following the eruption. The lowest point of the crater floor (justsouth of the center of the crater) is at an elevation of 2268 m relativeto datum and the perimeter of the floor is at an elevation b2350 m,whereas the surrounding pre-impact surface is at an elevation of~2360 m, meaning that the in-filling lava probably rose to the level ofhydrostatic equilibrium with eruptions in the region surrounding thecrater. The high point of the crater rim crest is 2662 m, giving a depth/

. See Fig. 2b for location. Spatter units are in red and rust-red, short spatter-fed flows inof spatter, and light green flows are prominent flows that appear to have been partiallyof the lava shield, and is comprised of undifferentiated flows where no continuous flowodel, are at 25-m intervals relative to the Mars datum. Mapped from THEMIS VIS imageswith the Central vent complex. See Fig. 2b for location. Color scheme of units is the samed, phase of undifferentiated flows at the perimeter of the lava shield. Contours, derivedars datum. Mapped from THEMIS VIS images at a spatial resolution of 18 m/pixel. (c,. 2b for location. Four separate vents can be identified in this area, and the overlappingts is the same as Fig. 3b. Contours, derived from the 128th-degreeMOLA digital elevationa spatial resolution of 18 m/pixel. (For interpretation of the references to colour in this

Fig. 4. Partially in-filled impact crater to the east of the central vent complex at 18.3°N, 246.0°E. See Fig. 2b for location. This crater appears to have been flooded by lava following theeruption of a fissure that intersected the western rim of the crater. At left we show the image of the crater, and at right the contours have been superimposed on the image. Thecontours (in white) are in increments of 10 m, and show that large areas of the central portion of the crater are almost flat. There has probably been subsidence of the central infillbecause the perimeter is ~60 m higher than the central part of the floor. Note the absence of lava flows or spatter on the rim of the crater, indicating the lack of high lava fountainsduring the eruption. At left and right, the image is a mosaic of THEMIS images V09678015, V10302011, and V10926044.


diameter ratio (i.e., 394/18,700) of 0.021. Typical freshMartian cratersof this diameter should have a depth of ~795 m according to thestudies of Boyce et al. (2006) or 605 m according to Garvin et al.(2003), so that it is reasonable to infer that there is ~200–400 m oflava infill within the crater.

Theeasternvent complex (17.9°N, 246.8°E) ismore complicated thanthe other two, in that here we find three overlapping lava shields andgreater diversity in the resultant landforms (Fig. 3c). Here the ventswere localized as individual craters with channelized flows emanatingfrom them. The western vent (vent #1) has a 3-km sinuous channelextending to the north. The middle of the three vents (vent #2) is a~1.9-km-wide lava lake feeding a 4.4-km channelized flow to thewest (Fig. 5a), and has a morphology that is very similar to the earlyphase of activity at Kupaianaha lava pond on Kilauea volcano, Hawai'i(Mangan et al., 1995; Kauahikaua et al., 1996) (Fig. 5b). Theeasternmost vent (vent #3) has a sinuous channel that superficiallyresembles a lunar sinuous rille. There is a 12.6-km-long fracture tothe southeast of vent #3, but no flows appear to have erupted fromthis site. The fracture appears to continue a further 14.4 km to thenortheast, but here the surface expression is a series of disconnected

Fig. 5. (a, left) Probable solidified lava lake within the eastern vent complex. See Fig. 2a for lothe Kupaianaha lava pond on Kilauea volcano (Fig. 5b). Mosaic of THEMIS images V139710appears to be a good analog to the solidified lava lake shown in Fig. 5a. The pond is activincandescent lava. The circular part of the pond in the foreground is ~100 m in diameter. U

pits, the easternmost of which has a raised rim. An isolated patch ofsmooth terrain, which is most likely spatter, appears to indicate thepresence of a fourth vent to the east of the other vents in thiscomplex. Very little lava erupted from this vent.

4. Thickness of flows

Mouginis-Mark and Christensen (2005) identified the fact that thelava flows within what we term the Central vent complex areunusually thin compared with other lava flows on Mars, and citedcases of flows only 4 m thick. In order to better estimate the thickness,and hence volume, of the flows, we have utilized the elevation dataprovided by individual MOLA ground tracks to obtain thicknessmeasurements of individual flows at each of the three vent complexes.Elevation data points on either side of the flow were used toextrapolate the pre-existing surface, and the average height of theflow surface above this pre-existing surface was estimated. Fig. 6illustrates the locations (all white circles and black squares) wherenear-orthogonal orientations of MOLA tracks and flow directionsallowed thickness measurements to be made. The mean thicknesses

cation. This is mapped as vent #2 in Fig. 3c. Note the similarity between this feature and10 and V18264018. (b, right) The Kupaianaha lava pond on Kilauea volcano, Hawai'i,e in this image, with the bright lines across the near-circular part of the pond being. S. Geological Survey photo.

Fig. 6. (a, left) Location of all spot elevation measurements made on the western vent complex. Black squares show locations of data that are correlated with measurements of flowwidth (Table 2) and the white circles showwhere additional data were collected. See Fig. 2a for location. Base image is THEMIS V05896014. (b, middle) Location of all spot elevationmeasurements made on the central vent complex. Symbols are the same as for Fig. 6 a, with the flow widths shown in Table 2. See Fig. 2a for location. Mosaic of THEMIS imagesV05484014, V11238004, V20161005, and V20473002. (c, right) Location of all spot elevation measurements made on vents 1 and 2 within the eastern vent complex. Symbols are thesame as for Fig. 6 a, with the flow widths shown in Table 2. See Fig. 2a for location. Mosaic of THEMIS images V13971010, V18264018, and V19200011.


for the Western, Central and Eastern vent complexes are 5.5±2.0 m(N=13), 4.3±1.4 m (N=57) and 4.3±1.7 m (N=11), respectively,and a t-test shows that these values are not significantly different atthe 70% level. The black squares in Fig. 6 identify depth measurementson flows that could be traced for relatively great (N5 km) distances, sothat representative width measurements could also be made, and thelength, width and depth data for these flows are summarized inTable 2, together with estimates of the corresponding flow volumes,V, obtained from the product of the measured length and the mean

Fig. 7. (a, left) Location of the ten photoclinometric profiles derived for the near-vent flowsdenote the lava levees identified in Fig. 7b. The openwhite circles denote the locations of theright) Ten individual topographic profiles across the lava channel identified in Fig. 7a. The lev~84×. Base image is THEMIS image V05484014.

width and depth. The mean depths of these selected flows, 5.6±2.2m, 4.6±0.7m and 4.3±1.7m for theWestern, Central and Easternvent complexes, respectively, are not significantly different from thoseof the bulk data set.

The thinness of these flows led us to make more detailedtopographic measurements of flows at two different distances fromthe vent. We used photoclinometry (a shape from shading techniquecommonly used to infer the pixel-scale topography of a scene from thevariation in brightness values in a calibrated image — the details are

on the southern rim of the central vent complex. See Fig. 2a for location. A–A′ and B–B′individual MOLA elevation measurements. Base image is THEMIS image V05484014. (b,ees of the channel are identified by the two lines A–A′ and B–B′. Vertical exaggeration is

Fig. 8. (a, left). Location of 12 photoclinometric profiles taken across a possible spatter rampart on the northern rim of the Central fissure. See Fig. 2a for location. The open circles arethe individual MOLA shots. The black circles for the top and bottom profiles denote the MOLA shots that are used to set the regional elevations (all profiles used these MOLA orbits,but not all tie points are shown). Base image is THEMIS image V05484014. (b, right). Twelve individual profiles across the spatter rampart on the northern rim of the central fissure.Shaded areas denote the spatter rampart. Profiles are presentedwith the northernmost profile at the top and the southern-most profile at the bottom. Vertical exaggeration is 122.5×.Examples of the definitions of the lava fountain parameters Λ and R used in section 7 are shown.

Fig. 9. (a, left) Locations of 30 photoclinometric profiles taken over a prominent lava flow to the north of the central vent complex. See Fig. 2a for location. The white arrows identifythe profiles illustrated in Fig. 9b. Base image is THEMIS frame V050484014. (b, right) Locations of 19 photoclinometric profiles taken across the prominent lava flow shown in Fig. 9a.The lava flow is shaded. Vertical exaggeration is ~430×.


Table 2Geometries of lava flows. Values given are the lengths, L, meanwidths,w, depths, d, andcorresponding total volumes, V, for selected lava flows from the Western, Central andEastern vent complexes (indicated by black squares in Fig. 6). The number ofmeasurements used to find the mean width and depth is shown in brackets after thevalue.

Flow number Length(km)

Mean width(m)

Mean depth(m)

Volume(106 m3)

(a) Western vent complex1 7.2 517 (7) 9.7 (1) 362 7.5 1120 (7) – (0) –

3 13.0 1030 (13) 3.9 (2) 524 12.5 960 (10) 5.6 (2) 675 10.2 770 (10) 4.9 (1) 39

(b) Central vent complex1 0 1570 (22) 5.1 (11) 2072 11.0 1040 (11) 2.8 (2) 323 33.0 2810 (28) 4.0 (5) 3744 20.0 1930 (9) 4.7 (5) 183

(c) Eastern vent complex1 (east vent #1) 21.0 1160 (18) 1.5 (1) 372 (east vent #1) 16.0 1310 (9) 6.0 (1) 1253 (east vent #1) 17.0 1430 (10) 5.0 (1) 1214 (east vent #2) 21.0 1200 (16) 4.5 (1) 1135 (east vent #2) 12.0 1450 (7) 3.0 (1) 526 (east vent #2) 16.0 1000 (9) 5.5 (1) 88

Table 4Rampart geometries. Values are given for the heights of the ramparts along the sides ofthe most active part of the fissure of the Central vent complex. See Fig. 8.

Outer northwest Inner northwest Southeast

Rampart height(m)

Rampart height(m)

Rampart height(m)

4.7 – –

4.3 – –

5.2 – –

2.3 0.6 0.92.2 2.2 0.43.8 2.0 0.83.1 1.1 1.64.0 0.7 1.64.0 1.5 1.92.0 1.1 1.02.5 1.3 0.62.0 0.3 0.6– – 1.4– – 1.5– – 2.0– – 1.0Means: 3.34 1.20 1.18

Table 5Geometry of the Central vent complex flow examined in detail. Values are given for thewidth, depth, cross sectional area, and volume increment as a function of distance fromthe vent. Photoclinometric profiles and locations shown in Fig. 9.

Profileno.

Distance fromflow front(m)

Flowwidth(m)

Flowdepth(m)

Cross-sectionalflow area(m2)

Volumeincrement(106 m3)

1 298 4010 5.90 23,700 7.062 597 3380 3.30 11,100 3.323 895 2530 2.60 6570 1.964 1190 2730 3.20 8720 2.605 1490 2960 4.00 11,800 3.536 1790 3030 5.20 15,800 4.717 2090 2780 3.50 9730 2.908 2390 2920 4.50 13,200 3.929 2680 1980 2.90 5740 1.7110 2980 1420 2.70 3820 1.1411 3280 1070 2.00 2140 0.6412 3580 1070 2.30 2460 0.7313 3880 1400 2.80 3920 1.1714 4180 1420 2.70 3820 1.14


given in Appendix 1) to study the topography of a lava channel thatextends 1 to 4.7 km from the rim of the apparently most active part ofthe fissure of the Central complex (Fig. 7a and b). This fissure providesthe best evidence for spatter ramparts on both sides of the fracture,implying that the eruption included a fire fountaining episode that feda number of rootless flows on both sides of the fissure by coalescenceof hot pyroclasts falling from the fountain, while at the same timeallowing some of the pyroclastic material to fall back to the surface at asufficiently low temperature that it remained as spatter close to thefissure. Table 3 shows that both levees stand about 2 m above the pre-existing surface and that the depth of the lava in the proximal part ofthis channel is just less than 1 meter. A photoclinometric analysis ofthe fissure ramparts (Fig. 8a and b) shows (Table 4) that the innerramparts are about 2 m high and the northwestern outer rampart is~3.3 m high. These values are very similar to the heights of the leveeson the drained channel, and we propose that these levees areessentially accumulating rampart material that was rafted away by asurge in the flow near the end of the eruptive activity, similar to whathas been observed in the ongoing eruptions at Kilauea, Hawai'i (Wolfeet al., 1988). Certainly the stratigraphic relationship between thedrained channel and the ramparts implies that they correspond to thelast two episodes of activity at the fissure, respectively.

We also investigated the variation of flow thickness and width as afunction of distance down-slope from the vent for one of the largerflows north of the fissure at the Central vent complex (Fig. 9) toestablish howmuch could be inferred about the flow dynamics. This issimilar to the analysis of much longer lava flows by Baloga et al.(2003), Glaze et al. (2003), Rowland et al. (2004), Glaze and Baloga

Table 3Measurements of levee heights and lava depth in channel for channel shown in Fig. 7ausing photoclinometric data in Fig. 7b.

Profilenumber

West levee East levee Central channel

Height (m) Height (m) Depth (m)

1 2.3 1.7 0.82 2.4 2.2 1.13 1.9 2.3 1.34 1.9 1.9 0.85 2.4 2.0 0.86 1.9 2.2 0.67 2.1 1.8 1.1Means 2.13 2.01 0.93

(2006, 2007), Hiesinger et al. (2007) and Garry et al. (2007), exceptthat here we employ photoclinometry (Appendix 1) to utilize thehigher spatial resolution of the THEMIS VIS instrument (18 m/pixel)rather than the 300-meter spacing of raw MOLA data. This 22-km-long flow starts 6 km from the rim of the fissure and is one of the mostprominent individual flows that we mapped at this complex (Fig. 3b).The distal part of this flow has a mean thickness of ~3.4 m and a meanwidth of 1870 m (see Table 5).

15 4470 1570 2.40 3750 1.1216 4770 1620 3.00 4850 1.4517 5070 1290 2.80 3610 1.0818 5370 1360 3.20 4360 1.3019 5667 870 2.00 1740 0.5220 5965 1780 3.80 6750 2.0121 6264 1800 4.00 7180 2.1422 6562 1540 3.00 4620 1.3823 6860 Complications due to presenceof impact crater 1.5024 7160 1450 3.90 5660 1.6925 7460 1200 4.30 5140 1.5326 7760 1140 4.90 5590 1.6727 8050 1400 3.30 4610 1.3728 8350 1250 3.00 3740 1.1129 8650 1250 3.00 3750 1.1230 8950 1960 3.10 6070 1.81

Mean: 1867 Mean: 3.36 Sum: 57.84

Table 6Geometry and yield strength of flows. Values are given for flow unit thicknesses, d, localsurface slopes, α, and lava yield strengths, y, as a function of down-flow distance fromthe vent. Data are derived from MOLA profiles alone.

d(m)

z(km)

sin α y(Pa)

West vent, flow no.1 9.7 6.8 0.0065 5882 4.5 4.6 0.0100 4143 6.6 7.1 0.0088 5424 4.9 3.9 0.0175 7955 4.0 2.0 0.0051 1886 5.0 4.2 0.0036 1667 4.3 4.2 0.0050 1998 9.4 6.8 0.0069 6069 6.2 7.1 0.0087 50710 3.5 7.4 0.0133 43711 3.8 13.6 0.0051 17812 5.4 15.6 0.0077 38513 3.7 17.0 0.0095 328

Mean: 410±196

Central vent, flow no.1 4.0 20.4 0.0010 362 4.6 19.4 0.0008 363 3.6 16.5 0.0014 474 3.9 16.0 0.0025 915 4.0 15.3 0.0013 476 3.0 13.1 0.0043 1207 5.7 14.0 0.0021 1098 3.9 14.2 0.0026 949 6.0 15.5 0.0020 10910 4.1 10.3 0.0053 20211 3.0 9.5 0.0027 7712 3.2 9.0 0.0042 12413 5.3 7.4 0.0041 20114 5.1 7.1 0.0066 31515 6.5 10.5 0.0028 16916 6.4 10.3 0.0042 24817 4.7 8.2 0.0044 19118 4.4 6.7 0.0054 22319 5.2 5.4 0.0034 16520 4.6 5.4 0.0090 38621 3.0 7.0 0.0022 6122 2.6 5.6 0.0074 17523 0.8 3.6 0.0035 2724 5.0 2.9 0.0084 39425 5.6 5.4 0.0073 38426 2.6 6.2 0.0048 11727 1.6 4.0 0.0065 9828 3.1 5.7 0.0060 17329 1.3 5.5 0.0032 4030 5.0 10.4 0.0017 8131 2.9 12.8 0.0047 12632 6.7 9.6 0.0056 35133 3.8 7.9 0.0058 20134 3.5 10.1 0.0067 21935 3.0 10.6 0.0036 9936 5.3 12.0 0.0034 16737 4.5 10.6 0.0036 14838 8.6 13.3 0.0032 25639 4.5 12.8 0.0038 15840 5.6 12.3 0.0051 26641 4.0 13.0 0.0034 12742 4.8 14.8 0.0031 14143 4.5 14.9 0.0042 17344 3.1 18.3 0.0051 14945 3.9 17.2 0.0017 6346 4.0 18.9 0.0036 13447 5.3 17.7 0.0043 21048 5.3 18.4 0.0022 10849 4.6 19.1 0.0036 154

Mean: 159±93

East vent 1, flow no.1 1.5 6.9 0.0092 1292 6.0 14.2 0.0085 475

(continued on next page)

Table 6 (continued)

d(m)

z(km)

sin α y(Pa)

East vent 1, flow no.3 5.0 15.3 0.0051 2364 3.5 18.7 0.0045 145

Mean: 246±160East vent 2, flow no.1 6.0 11.9 0.0057 3162 4.5 8.7 0.0050 2083 6.5 7.5 0.0050 3004 3.0 1.6 0.0042 1165 2.4 3.5 0.0068 1526 3.0 2.4 0.0038 1077 5.5 6.0 0.0037 187

Mean: 198±83


5. Analysis offlowmorphology, rheologyandemplacementdynamics

5.1. Morphology and rheology

None of the lava flow units from any of the three vent complexesshow any clear evidence of a levee-and-central channel structureexcept very near their vents (e.g. Fig. 7). At distances beyond ~5 kmthe flows generally have a slightly concave upward profile (e.g.Fig. 9b) that shows no systematic changewith distance along the flow.We therefore assume that in general these lavas propagated as sheetflows, so that we can define a mean flow speed u(z) as well as themean flow depth d(z) at any distance z from the vent (see Notationsection for definitions of all variables). We further assume that thelavas behaved essentially as Newtonian fluids. Clearly this is notstrictly possible, as a Newtonian fluid would have continued to spreadlaterally as well as down-slope from the vent except where restrictedby pre-existing topography and we do not see that topography is asignificant factor controlling the widths of most flow units. Inevitablycooling of material at the front and margins of any flow must occur,and this can induce non-Newtonian rheology, specifically an apparentyield strength that allows levees, which must be very narrow in thepresent examples, to form and limit spreading of the flow.

We estimated the apparent yield strength of flow units in the usualway by equating the gravitational shear stress acting down thesubstrate slope, α, to the yield strength, y, i.e.:

y = ρ g d sin αð Þ ð1Þ

where ρ is the lava density (assumed to be 2500 kg m−3 to allow forsome possible vesicularity of the lava) and g is the acceleration due togravity, ~3.72 m s−2. The flow depths, d, needed for this calculationwere the 81 values based on MOLA profiles indicated in Fig. 6. Theresulting values of y are shown in Table 6 and are plotted as a functionof distance from the vent in Fig. 10a–d. The values are very scattered,but seem systematically slightly smaller for the Central vent than forthe Western vent or either of Eastern vents 1 and 2. The scatter isunderstandable in that local changes of substrate slope, smalltopographic obstacles and fluctuations in volume flux from the ventmay all act to increase the depth of a lava flow locally, resulting in anincrease in the yield strength inferred from the above equation,whereas once a yield strength is established, no changes in conditionscan cause the flow depth to decrease below the value set by that yieldstrength. This implies that the most representative measure of theapparent yield strength that can be obtained from Fig. 10 is the lowerenvelope of the values. Possible values and trends of this lowerenvelope are shown in Fig. 10; there is a hint of an increase in y withdistance from the vent, something that might be anticipated becausethe amount of lava cooling must increase with distance traveled. All ofthe lower envelope values (~25–100 Pa) are much smaller than yieldstrengths associated with mafic lava flows on Earth, typically 103 to

Fig.10. The variation of inferred lava yield strength, y, with distance from the vent, z, for the vents identified in Figs. 2 and 3. In each case two possible trends for the lower envelope ofthe values are shown; see text for discussion.


104 Pa (see summary in Wilson and Head, 1994), though a value assmall as 350 Pa was reported for a flow at Mauna Loa, Hawai'i (Moore,1987). This result is consistent with the absence of clearly-definedlevees on these martian flows. We examine the possible viscosity ofthe lava in the next section as part of the analysis of the flow dynamics.

5.2. Flow dynamics

With the assumption of Newtonian rheology, the mean speed, u, ofa flow can be related to the flow depth, d, the density, ρ, and viscosity,η, of the lava, the acceleration due to gravity, g, and the slope of thepre-existing topography, α, by standard fluid-dynamics equations(e.g. Knudsen and Katz, 1958), either:

u = ρ g d2 sinα� �

= 3ηð Þ ð2aÞ

if the flow motion is laminar, or:

u = 2 g d sinαð Þ = f½ �1=2 ð2bÞ

if the flow motion is turbulent, where f is a friction factor. The frictionfactor varies with the amount of turbulence in the flow, characterizedby the Reynolds number for the motion, Re, given by:

Re = 4 d u ρð Þ= η ð3Þ

In fully turbulent flow f can be approximated by:

f = 0:32 = Re0:25 ð4Þ

(Huppert and Sparks, 1985), whereas in fully laminar flow:

f = 24 = Re ð5Þ

The choice of which of Eqs. (2a) and (2b) to use for the flow speedis dictated by evaluating Re using the speed from each equation inturn and choosing whichever one yields a consistent value, i.e. a valueof Re N~ 2000 if the flow is turbulent or bb~ 2000 if the flow islaminar. As pointed out by Wilson and Head (1981), empirically allthat is needed is to evaluate both expressions for u and take thesmaller speed. This action always satisfies the Reynolds numbercriterion; it introduces a small error of ~10% in the speed for flows inthe middle of the transition region between fully laminar and fullyturbulent flow, but this error is negligible relative to the accumulatedeffects of the errors in measurements of d and α.

A major problem with attempting to deduce the dynamics ofany lava flow from morphological measurements alone is that wedo not know the viscosity of the lava. If the lava motion is fullyturbulent, Eq. (2b) shows that the flow speed does not depend onthe viscosity, but the Reynolds number marking the transition fromlaminar to turbulent flow does involve the viscosity, and so wecannot be sure whether a flow was laminar or turbulent unless weknow the viscosity. However, we can at least determine the valueof the viscosity, ηt, flow speed ut, Reynolds number, Ret, andfriction factor, ft, marking the transition between the two flowregimes. If we equate Eqs. (4) and (5), we have Ret

3/4=75, so thetransition must occur close to the value Ret=316; either of Eqs.(4) or (5) then show that the corresponding value of ft must be0.0759. At this transition point between flow regimes, Eqs. (2a)

Table 7Properties of flows listed in Table 2. Values are given for widths, w, lava depths, d, andground slopes, α, together with the flow speeds, ut, volume fluxes, Qt, andemplacement times, τ, if the flows were turbulent, and the maximum lava viscosity,ηt, that would have allowed turbulence to occur.

Flowno.

z(km)

w(m)

d(m)

sin α ut(m s−1)

ηt(Pa s)

Qt

(m3 s−1)τ(hours)

(a) Western vent complex1 7.2 612 9.7 0.00653 2.49 763 14,770 0.83 4.0 864 4.3 0.00501 1.45 198 5400 2.53 7.0 828 3.5 0.01328 2.14 236 6190 1.74 4.5 684 5.0 0.00358 1.33 209 4530 2.64 7.5 918 6.2 0.00875 2.31 452 13,100 1.55 4.2 846 4.9 0.01745 2.90 449 12,000 1.0

(b) Central vent complex1 4.5 970 4.60 0.00902 2.02 293 9000 3.61 5.5 970 5.19 0.00342 1.32 216 6640 5.51 6.0 1060 4.40 0.00545 1.53 213 7120 4.71 8.0 990 4.68 0.00439 1.42 210 6600 5.11 9.0 1780 4.40 0.01374 2.43 339 19,100 3.01 10.0 1850 5.25 0.00274 1.19 197 11,600 6.11 11.0 1310 6.38 0.00418 1.62 326 13,600 4.51 12.0 1500 6.49 0.0028 1.33 274 13,000 5.41 17.0 2240 5.45 0.00244 1.14 197 13,900 6.31 18.0 2470 5.64 0.00315 1.32 235 18,400 5.51 21.0 2160 3.15 0.00223 0.83 83 5630 8.72 6.0 1000 2.55 0.00739 1.36 110 3480 2.22 8.0 1530 3.00 0.00219 0.80 76 3680 3.83 8.0 3650 4.720 0.00277 1.13 169 19,500 8.13 9.0 3640 5.000 0.00173 0.92 146 16,800 10.03 12.0 2290 2.910 0.00467 1.15 106 7700 7.93 13.0 1690 3.280 0.00467 1.23 127 6800 7.53 17.0 2320 4.260 0.00156 0.81 109 8000 11.44 15.0 1880 4.48 0.00415 1.35 191 11,400 4.14 17.0 2050 3.94 0.00173 0.82 102 6600 6.84 18.0 2090 5.30 0.00426 1.49 249 16,500 3.74 19.0 2090 5.27 0.00220 1.07 178 11,700 5.24 20.0 3000 4.64 0.00358 1.28 187 17,800 4.4

(c) Eastern vent complex1 (vent #1) 6.0 1160 1.5 0.00924 1.17 55 2030 5.02 (vent #1) 13.0 1960 6.0 0.00851 2.24 424 26,300 2.03 (vent #1) 16.0 1480 5.0 0.00508 1.58 249 11,700 3.04 (vent #2) 15.0 1380 4.5 0.00496 1.48 210 9200 4.05 (vent #2) 11.0 2390 3.0 0.00416 1.11 105 7900 3.06 (vent #2) 13.0 1310 5.5 0.00366 1.40 244 10,200 3.2

Table 8Data for each of the profiles across the Central vent flow examined in detail (Table 5 andFig. 9). Values are given for flow widths, w, lava depths, d, and ground slopes, α,together with the flow speeds, ut, volume fluxes, Qt, and emplacement times, τ, if theflows were turbulent, and the maximum lava viscosity, ηt, that would have allowedturbulence to occur.

Profile z(km)

w(m)

d(m)

sin α ut(m s−1)

ηt(Pa s)

Qt

(m3 s−1)τ(hours)

1 300 4010 5.90 0.00149 0.93 173 22,000 6.62 600 3380 3.30 0.00154 0.71 74 7870 8.73 900 2530 2.60 0.00159 0.64 52 4190 9.64 1190 2730 3.20 0.00165 0.72 73 6270 8.55 1490 2960 4.00 0.00170 0.82 103 9670 7.56 1790 3030 5.20 0.00175 0.95 155 14,900 6.57 2090 2780 3.50 0.00181 0.79 87 7660 7.88 2390 2920 4.50 0.00186 0.91 129 11,900 6.89 2680 1980 2.90 0.00191 0.74 68 4230 8.310 2980 1420 2.70 0.00197 0.72 62 2760 8.511 3280 1070 2.00 0.00202 0.63 40 1350 9.712 3580 1070 2.30 0.00207 0.68 50 1680 8.913 3880 1400 2.80 0.00213 0.76 68 3000 8.014 4180 1420 2.70 0.00218 0.76 65 2900 8.015 4470 1560 2.40 0.00223 0.72 55 2720 8.416 4770 1620 3.00 0.00229 0.82 78 3980 7.517 5070 1290 2.80 0.00234 0.80 71 2900 7.618 5370 1360 3.20 0.00239 0.87 88 3780 7.119 5670 870 2.00 0.00245 0.69 44 1206 8.820 5970 1780 3.80 0.00250 0.96 116 6520 6.321 6260 1800 4.00 0.00255 1.00 127 7180 6.122 6560 1540 3.00 0.00260 0.88 83 4050 7.024 7160 1450 3.90 0.00271 1.02 126 5760 6.025 7460 1200 4.30 0.00276 1.08 147 5550 5.726 7760 1140 4.90 0.00282 1.16 180 6500 5.327 8050 1400 3.30 0.00287 0.96 101 4438 6.328 8350 1250 3.00 0.00292 0.93 88 3460 6.629 8650 1250 3.00 0.00298 0.94 89 3510 6.530 8950 1960 3.10 0.00303 0.96 94 5830 6.4Mean values: 1870 3.36 0.00226 0.85 93 5780 7.4


and (2b) must yield the same velocity, and equating them andsimplifying we have:

ρ2g d3 sinα ft = 18η2t ð6Þ

Inserting ft=0.0759 and rearranging we find:

ηt = 0:0649 ρ2g d3 sinα� �1=2 ð7Þ

Finally from Eq. (2a) we then have, at the transition:

ut = 5:13 g d sinαð Þ1=2 ð8Þ

Using the same individual flows from the Western, Central andEast vent complexes shown in Table 2, the above analysis was used todetermine the values of flow speed, ut, and viscosity, ηt, marking thetransition between laminar and turbulent flow regimes for alllocations where a flow depth and ground slope could be measured.We emphasize that, because the flow speed does not depend on thelava viscosity in turbulent flow, this value of ut is the maximum speedthat the flow could have had. Thus multiplying ut by the width anddepth of the flow yields a maximum estimate of the volume flux fromthe vent feeding the flow, Qt. Then dividing the volume flux by the

flow volume, V, yields the eruption duration, τ. The values are given inTable 7. Clearly, all of these flowswould have been turbulent if the lavaviscosity were less than some value in the range ~50–200 Pa s; thecorresponding maximum flow speeds would have been in the range0.8 to 2.5 m s−1; the maximum volume eruption rates would havebeenwithin a factor of about 2 of 10,000m3 s−1, and the emplacementtimes span the range ~1 to ~10 h.

The above analysis was also carried out for the single flow to the NEof the Central vent complex (Table 5 and Fig. 9). The results, given inTable 8, are similar to those for the flows for which we have fewerdata: this flow would have been turbulent if the lava viscosity hadbeen less than ~100 Pa s; it would have had a maximum mean flowspeed of ~0.85 m s−1 and a volume flux of ~5780 m3 s−1, with anemplacement time of 7 to 8 h.

The above treatment yields the maximum speeds and volumefluxes of flows. A lower limit on the eruption rate, and a constraint onthe maximum lava viscosity, can be obtained by noting that thelengths of lava flow units are determined in one of two ways. A flowunit can be volume-limited or cooling-limited (Guest et al., 1987). Aflow unit is volume-limited if its length stops increasing whenmagmaceases to be erupted at the vent. Some small amount of drainage of thehottest, central part of the flow may occur, but a distinct flow front isnormally visible at the point where drainage begins. Conversely a flowunit is cooling-limited if its length stops increasing as a result of theexcessive cooling of the lava. It is unlikely that this condition willhappen to coincide with the cessation of lava eruption from the vent,and so to accommodate the continuing supply, either a new flow unitmust be formed at the vent, or a breakout must occur from some pointon the margin or front of the original flow unit. If the breakout occursat the flow front, its width and thickness will normally readilydistinguish it from a drainage feature at the front of a volume-limitedflow unit. If a flow unit is cooling-limited, its volume eruption rate can

Table 10Parameters for the profiles across the Central vent complex flow studied in detail(Table 5 and Fig. 9). Values are given for the volume flux from the vent, V, the mean flowspeed, u, the implied lava viscosity, η, and the emplacement time, τ, derived on theassumption that the length was limited by cooling rather than cessation of the eruptionat the vent.

Profile no. V(m3 s−1)

u(m s−1)

η(Pa s)

τ(days)

1 786 0.033 4840 2.12 1182 0.106 491 1.43 1123 0.171 196 1.54 984 0.113 464 1.75 854 0.072 1170 1.96 674 0.043 3440 2.47 917 0.094 728 1.88 750 0.057 2050 2.29 788 0.137 363 2.110 606 0.158 281 2.711 619 0.289 87 2.712 537 0.218 156 3.113 578 0.147 351 2.814 606 0.158 311 2.715 751 0.201 199 2.216 622 0.128 497 2.617 532 0.147 386 3.118 492 0.113 673 3.319 503 0.289 105 3.320 540 0.080 1400 3.021 518 0.072 1750 3.222 593 0.128 566 2.824 430 0.076 1680 3.825 321 0.062 2540 5.126 269 0.048 4360 6.127 488 0.106 914 3.428 479 0.128 636 3.429 481 0.128 647 3.430 730 0.120 751 2.3

Mean values: 647 0.125 1100 2.7


be deduced from its length. However, if it is assumed that a given flowunit was cooling-limited, whereas it was in fact volume-limited, thenthe volume eruption rate deduced from its length will alwaysrepresent a lower limit on the actual volume eruption rate feedingthe flow. The relationship between predicted volume effusion rate Qc

and observed flow length L for cooling-limited flows is discussed byPinkerton and Wilson (1994) who show that:

Qc = L κ Gzwð Þ= β dð Þ ð9Þ

where w and d are the width and depth of the flow as before, κ is thethermal diffusivity of the lava (generally close to 7×10−7 m2 s−1), Gz isa critical value of the dimensionless Grätz number, equal to about 300,and β is a factor that depends on the ratio of the effective hydraulicradius of the flow to its actual depth. For flows that are coming to restdue to cooling, the cooled surface crust will not be moving as fast as thehotter lava in the core of the flow and the relevant value of β is 4.

Using the data on flow lengths and depths in Table 2, Eq. (9) wasused to generate the cooling-limited volume fluxes listed in Table 9.Values cover a wide range, from 20 to 1200 m3 s−1. All of the impliedflow conditions are laminar, with flow speeds generally within a factorof 2 of 0.05 m s−1. Implied viscosities cover a very wide range, from~100 to nearly 5×105 Pa s (though clearly the largest of these values isvery improbable for a mafic lava), and emplacement times aregenerally a few days. The corresponding analysis was carried out forthe flow studied in more detail (Table 5 and Fig. 9), and Table 10 givesthe resulting parameter values which, not surprisingly, are much lessscattered. The implied volume flux is ~650 m3 s−1 and the mean flowspeed is ~0.13 m s−1. The flow motion is laminar, with an impliedviscosity of ~1100 Pa s; the emplacement time is 2.7 days.

We now have two extreme sets of estimates for these flows: anupper limit on the flow speed and volume eruption rate if the flowmotion was turbulent, and a lower limit on these quantities if themotionwas laminar and the flow lengths were limited by cooling. Theupper limit values require the lava viscosity to have been less thanabout 100 Pa s, whereas the lower limit values require generally muchgreater viscosities — in many cases values much larger than any thatwould normally be associated with mafic lavas. Such large viscositiesdo not seem consistent with the extreme thinness of these flows —

recall that the implied yield strengths were only of order 100 Pa. Alsowe note that the levees between the proximal parts of the flows on the

Table 9Parameters for flows listed in Table 2. Values are given for the volume flux from the vent,V, the mean flow speed, u, the implied lava viscosity, η, and the emplacement time, τ,on the assumption that flow lengths were limited by cooling rather than cessation of theeruption at the vent.

Flow no. V(m3 s−1)

u(m s−1)

η(Pa s)

τ(days)

(a) Western vent complex1 24 0.004 472,000 17.52 No data available3 180 0.045 9610 3.34 112 0.021 28,600 6.95 85 0.022 58,200 5.3

(b) Central vent complex1 424 0.053 7230 5.62 217 0.075 1530 1.73 1210 0.106 1460 3.64 429 0.047 4690 4.9

(c) Eastern vent complex1 (east vent #1) 853 0.490 132 0.52 (east vent #1) 275 0.023 40,700 5.33 (east vent #1) 263 0.036 11,000 5.34 (east vent #2) 337 0.054 5720 3.95 (east vent #2) 502 0.070 1660 1.26 (east vent #2) 201 0.028 12,400 5.1

south side of the Central fissure vent visible in Figs. 6b and 7a have“streamlined” shapes, strongly suggests scouring or sculpting of theselevees by fast moving lava. We therefore consider it extremely likelythat the actual flow conditions were close to or in the turbulentregime, with the lava viscosity being ~100 Pa s, the volume eruptionrates being in the range 5000 to 10,000 m3 s−1, the flow speeds being1 to 2 m s−1 and the emplacement times being typically 5 h.

6. Magma sources and the rise of magma to the surface

6.1. Magma source depths

All magma reaching a planetary surface is likely to have itsultimate source somewhere in the mantle. However, it is wellestablished that magma rising from great depth may be trapped invarious ways to form reservoirs in which chemical evolution takesplace. A common method of trapping is for melts to rise buoyantlyfrom their source zone through host rocks of decreasing density untilthe melts reach a level of neutral buoyancy (Ryan,1987, 1988). Neutralbuoyancy reservoirs may be located at the base of the crust on Earth(Andrew and Gudmundsson, 2008) or within growing volcanicedifices on any planet (Head and Wilson, 1992; Wilson and Head,1994). In the present case the fissure vents are not associated withmajor volcanic edifices like the Tharsis and Elysium shield volcanoes.Also, we see neither collapse calderas, the presence of which isgenerally taken as evidence of the presence of a shallow reservoir in avolcano (Folch and Marti, 2004), nor patterns of surface faulting dueto stresses imposed by reservoir inflation and deflation, such as wasinferred for Olympus Mons by Zuber and Mouginis-Mark (1992).However, this (lack of) evidence does not preclude the possibility thatthe small constructs studied here were fed from crustal reservoirs toodeep, or too small, to produce surface expressions. Another

Table 11Implied crustal properties. For a range of values of the surface void space fraction vs,values are given for the density ρo in kgm-3 of the rocks at the base of the crust requiredto produce a mean crust density of 2950 kg m-3, and for the magma neutral buoyancydepth for two cases: volatile-free magma and magma containing 0.3 mass % CO2 in itsmantle source region. The ambient lithostatic pressures in the crust host rocks are alsogiven. A null entry means that no neutral buoyancy level exists and magma is buoyantall the way from the base of the crust to the surface.

vs ρ0 Neutral buoyancy depth in km for: Corresponding lithostatic pressure inMPa at neutral buoyancy depth for:

Gas-free magma 0.3 mass % CO2 Gas-free magma 0.3 mass % CO2

0.050 2970 – – – –

0.060 2980 1.2 – 12.6 –

0.070 2980 2.2 – 23.1 –

0.075 2990 2.7 – 28.0 –

0.100 3000 4.5 – 46.1 –

0.150 3030 6.8 – 69.0 –

0.160 3030 7.2 – 72.6 –

0.170 3040 7.6 3.4 75.8 32.60.175 3040 7.7 4.6 77.2 45.20.200 3050 8.4 6.2 83.1 60.70.238 3080 9.3 7.7 90.5 74.30.250 3080 9.6 8.1 92.8 77.40.300 3120 10.5 9.3 99.2 87.30.313 3120 10.6 9.6 100.4 89.20.400 3190 11.8 11.0 106.9 98.60.500 3270 12.9 12.3 110.3 104.0


consideration is the fact that neutral-buoyancy-level reservoirs canfeed not only upward propagating dikes but also dikes that intrudelaterally, centered on the neutral buoyancy level (Ernst et al., 1995,2001). If the reservoir is deep, such dikes can extend for very greatdistances before they reach the surface and erupt. However, havingexamined the regions along strike from each of the fissure vents wefind no evidence for possible sources of any of the dikes; nor do wefind examples of the same dike cropping out to form more than onevent. These various lines of evidence suggest that the fissure ventsstudied here were fed by dikes propagating mainly vertically fromreservoirs located at great enough depths that no surface expressionof the reservoir is visible. In an attempt to better constrain thesedepths, and to explore their influence on the magma eruption rate, wenow examine two scenarios: reservoirs within the crust located atlevels of neutral buoyancy, and reservoirs at the base of the crust.

6.2. Intra-crustal reservoirs

To establish the possible positions within the crust of neutralbuoyancy level magma reservoirs we need values for the densities ofthe crust and of the magmas being erupted. The properties of themartian lithosphere have been explored by several groups usinginversion of gravity and topography data. Zuber et al. (2000) derived amean crustal density of 2900 kg m−3 and inferred a crustal thicknessof 50 km in the Tharsis region. Turcotte et al. (2001) derived a meancrustal density of 2950 kg m−3. McKenzie et al. (2002) found a meancrustal density very close to 3000 kg m−3, with the correspondingcompact density given as 3320 kg m−3 and the mantle density as3500 kg m−3. They estimated an elastic lithosphere thickness inTharsis of ~70 km but did not given an explicit crustal thicknessestimate. We adopt 2950 kg m−3 as the mean crustal density and50 km as the crustal thickness. Scaling from the values cited byMcKenzie et al. (2002), this would imply a compact crustal basaltdensity of 3265 kg m−3. Boudreau and Philpotts (2002) show that theratio of the densities of a typical solid terrestrial basalt at a temperatureof 800 °C and the liquid fromwhich it cools is 1.155; thus a compact soliddensity of 3265 kg m−3 would correspond to a liquid density of2826 kgm−3, almost exactly in themiddle of the range 2800 to 2850 kgm−3 derived by McSween (2002) from considering the density ofmartianmeteorites. We adopt 2825 kgm−3 as the basalt liquid density.

Based on the pore space compaction profiles for terrestrial volcanicregions,HeadandWilson(1992)proposedageneralmodel for thedensitystructure within volcanic constructs and regions of volcanically emplacedcrust. The density ρc is given as a function of the depth z by:

ρc = ρo = 1 + vs = 1− vsð Þf g exp −λρogzð Þ½ � ð10Þ

where ρ0 is the density at great depth where all of the pore space hasbeen compressed to zero, vs is the pore space volume fraction at thesurface, and λ is a gravity-independent compaction parameter equalto 1.18×10−8 Pa−1. The resulting variation of pressure with depth is:

P = λ−1ln vs + 1− vsð Þexp λρogzð Þ½ � ð11Þ

It is observed that terrestrial volcanic constructs have surfacedensities of ~2200 kg m−3 (Hill, 1969; Gudmundsson, 1987),corresponding to vs=~0.24; Wilson and Head (1994) pointed outthat the low martian atmospheric pressure would encourage volatilerelease frommartian magmas. Depending onwhether gas bubbles areretained in erupted rocks or lost in explosive activity, the density ofsurface volcanic deposits on Mars could be either less or greater thanon Earth, and so we make no presupposition about the value of vs, butexplore a range of values. For each choice of vs we use a range of valuesof ρ0 to calculate the variation of crustal density with depth andobtain the mean value down to the crust thickness of 50 km. We thusfind the value of ρ0 that yields a mean crust density of 2950 kg m−3.

Table 11 shows the results: ρ0 is required to range from 2974 to3270 kg m−3 as vs varies from 0.05 to 0.5. We next seek the depth inthe crust at which the local density is equal to the magma density,2825 kg m−3, and identify this as the likely depth of the center of anymagma reservoir. These depths are shown in Table 11. If the value of vsis smaller than ~0.05, no neutral density level exists and no magmareservoir is expected to form. At vs=0.06 a reservoir could form at adepth of ~1.2 km, and as vs increases the depth of the center of thepotential reservoir increases, reaching ~9.6 km at vs=0.25 (similar tothe terrestrial value) and ~12.9 km if vs is as large as 0.5.

There is a complication neglected in the above treatment. Mantlemagmas on Earth contain significant amount of volatiles, mainly CO2,H2O with smaller amounts of S species, and we assume the same istrue on Mars. In a study of Kilauea volcano on Earth, Gerlach (1986)estimated the amounts of these volatiles as 0.3mass% H2O, 0.65mass%CO2 and 0.13 mass% S. The solubilities of CO2 and H2O as a function ofpressure, P, in mafic magmas are given by Dixon (1997) as:

nCO2= 5:9 × 10−12P + 5:0 × 10−6 ð12Þ

and

nH20 = 6:8 × 10−8P0:7 ð13Þ

where the values of n are mass fractions and P is expressed in Pascals.Examination of the crustal pressures implied by Eq. (11) readily showsthat ifwaterwerepresent in the amount implied for Kilaueanonewouldhave exsolved at any of the magma reservoir depths just found, but CO2

would. Indeed, CO2 would have started to exsolve from the magma bythe time it reached the base of the crust. We have therefore calculatedthe bulk density of a magma with a liquid density of 2550 kg m−3

exsolving CO2 as it rises through the crust in equilibrium with thelithostatic pressure of the host rocks, and found the depth at which thismagmawouldbeneutrally buoyant in eachof the crustaldensitymodelsdescribed above. A conservative initial CO2 content of 0.3 mass %, abouthalf of the estimated terrestrial value of Gerlach (1986), was assumed.Table 11 again shows the resulting potential magma reservoir centerdepths. In this case no reservoir can exist unless vs is at least 0.17, and theshallowest possible reservoir center depth is ~3.4 km.

The decision as to which of these two sets of potential intra-crustalreservoir depths is more likely to approximate the real case is not easy.The earliest batches of magma to be produced by the onset of mantle

Table 12Magma requirements to allow eruptions. For magma reservoirs at the base of the crust,values are given for the liquid magma density, ρ; the solid magma density ρcs; the meanvoid space fraction in the crust, v; the excess magma pressure available to drive eruptions,ΔP; the corresponding pressure gradients driving eruptions, dP/dz; the minimum depthbelow the base of the crust of amantlemagma source thatwill just raise a static column ofmagma to the surface,M; and the depth below the base of the crust of a magma source inthe mantle that will provide a pressure gradient of 500 Pa m-1,M'.

ρ(kg m−3)

ρcs(kg m−3)

v ΔP(MPa)

dP/dz(Pa m−1)

M(km)

M′

(km)

2553 2950 0.000 73.8 14802600 3000 0.018 65.1 13002650 3060 0.036 55.8 11202700 3120 0.054 46.5 9302750 3180 0.071 37.2 7442800 3230 0.088 27.9 5582825 3260 0.096 23.3 465 0.92850 3290 0.104 18.6 372 3.32900 3350 0.119 9.3 186 9.12950 3410 0.134 0.0 0 0.0 16.23000 3470 0.149 5.0 25.23050 3520 0.163 11.1 37.13100 3580 0.176 18.8 53.53150 3640 0.189 28.6 77.6


melting will not instantly create a reservoir on reaching a neutralbuoyancy level; very many discrete intrusions must occur, closelyspaced enough in time that one intrusion does not completely coolagainst its host rocks before another arrives (Wilson et al., 2001). Itmay be that the intense fracturing and deformation of the crust inthese early stages allows exsolved CO2 to escape easily from theintruded magma, so that ultimately it is the gas-vesicle-free densitythat determines the location of the reservoir, but this cannot beguaranteed. To explore the range of possibilities we consider bothscenarios, and so Table 11 shows values of the various parameterswhen vs=0.175, in which case a magma reservoir forms centered on7.7 km depth when there is no influence of magmatic CO2 in formingthe reservoir, and vs=0.238, in which case a magma reservoir formscentered on 7.7 km depth when magmatic CO2 does play a part indetermining the depth.

First consider the no-CO2 case. Table 11 shows that the host rockpressure at the 7.7 km depth level of the reservoir center is 77.2 MPa.Comparison of the computed densities of the magma, obtained using,e.g., Eq. (6) in Scott et al. (2002), and the host rock, obtained fromEq. (18), shows that the magma, which has been positively buoyanteverywhere below the neutral buoyancy level, continues to bepositively buoyant until a depth of 4.6 km at which depth it becomesnegatively buoyant. It remains negatively buoyant until it reaches adepth of 2.7 km below the surface, at which depth it becomespositively buoyant again all the way to the surface. The negativebuoyancy between the 4.6 and 2.7 km depths is more thancompensated by the positive buoyancy between the 7.7 and 6.6 kmdepths, and so no excess pressure is needed in the magma reservoir toallow this magma to reach the surface once a dike is initiated in thereservoir roof. The pressure corresponding to the static weight of themagma column from the reservoir center to the surface is 70.8 MPa,and so the difference between the 77.2 MPa reservoir pressure andthis weight, i.e. 6.4 MPa, is available to drive the magma motionagainst the fluid friction with the dike walls. Dividing this pressuredifference by the path length of 7.7 km gives a motion-drivingpressure gradient, dP/dz, of 826 Pa m−1.

Next consider the alternative case where CO2 in the magma doesinfluence the reservoir location Table 11 shows that the host rockpressure at the 7.7 km depth level of the reservoir center is 74.3 MPa.Comparison of the computed densities of the magma and host rockshow that the magma is negatively buoyant from this level to a depthof ~1.3 km below the surface, and that an excess pressure of 1.6 MPa isneeded in the reservoir to raise the magma to this depth, requiring areservoir pressure of 75.9 MPa. Once raised to the 1.3 km depth themagma is positively buoyant to the surface and so is able to erupt. Thepressure corresponding to the static weight of the magma columnfrom the reservoir center to the surface in this case is 70.5 MPa, and sothe difference between the 75.9 MPa reservoir pressure and thisweight, ~5.4 MPa, is available to drive the magma motion against thefluid friction with the dike walls. Dividing this pressure difference bythe path length of 7.7 km gives a motion-driving pressure gradient ofdP/dz=698 Pa m−1.

The two examples just given can be repeated for a range of otherpotential reservoir depths. If the reservoir center depth is at ~9.6 km,the pressure gradients corresponding to the above examples are 520and 490 Pa m−1; for a reservoir depth of ~11.8 km the pressuregradients are 350 and 324 Pa m−1. Clearly, as the reservoir depth getsgreater the available excess pressure to drive magma motiondecreases; but, within a factor of ~2, a value of ~500 Pa m−1 ischaracteristic of the range found for reservoirs in the upper 10–15 kmof the crust. We now consider an alternative scenario.

6.3. Sub-crustal reservoirs

Here we presuppose that the magma reservoir for the eruptions isa sill-like accumulation of mantle melt at the base of the crust. If we

continue to use 2950 kg m−3 for the mean crustal density, ρcm, and2825 kg m−3 for the melt density, ρ, the melt will on average bepositively buoyant in the crust and so no excess pressure derived fromthe buoyancy of the melt in its mantle source is required to allow it toreach the surface. The excess of the lithostatic pressure at the base ofthe crust over the static weight of the magma in a dike extending tothe surface is ΔP given by:

ΔP = g C ρcm − ρð Þ ð14Þ

where C is the thickness of the crust. The pressure gradient, dP/dz,available to drive magma upward against wall friction in a dike is ΔP/C, i.e. [g (ρcm−ρ)], which is independent of the crustal thickness andhas the value 465 Pa m−1. We can explore the effects on the pressuregradient of the magma density being outside the 2800–2850 kg m−3

range suggested by McSween (2002) by supposing that the crust inthis part of Tharsis consists largely of solidified lavas of similarcomposition to those most recently erupted. We take the ratio of themelt density ρ to the solidified magma density ρcs to be the valuefound by Boudreau and Philpotts (2002), i.e. ρcs=1.155 ρ. If theaverage porosity (void space fraction) in the crustal rocks is v, therelationship between ρcs, ρcm and v is:

v = 1− ρcm = ρcs½ � ð15Þ

Table 12 shows how ρcs, v, ΔP and dP/dz vary with the assumedmagma density ρ. The smallest melt density consistent with a non-negative void space fraction is 2553 kg m−3. The largest melt densityallowing the magma to be positively buoyant in the crust is, of course,2950 kg m−3, the same as the mean crustal density. For densitiesbetween these values, the pressure gradient available to drive magmato the surface ranges from zero to ~1480 Pa m−1. The average of thesevalues, 715 Pa m−1, is not very different from the typical values foundin the previous section for intra-crustal magma reservoirs. Of course, ifthe magma is denser than the mean crustal density, the “excess”pressure ΔP available to drive the magma motion is negative, anderuptions cannot occur unless there is a super-lithostatic pressure inthe magma reservoir. Such an excess pressure would be present if thereservoir at the base of the crust had a continuous pressure connectionwith a mantle melt source. If the average mantle rock density is ρm,which McKenzie et al. (2002) suggest is ~3500 kg m−3, the minimum


depth of that melt source, M, must satisfy the relationship [g C(ρ−ρcm)]=[g M (ρm−ρ)], i.e.:

M = C ρ − ρcmð Þ= ρm − ρð Þ½ � ð16Þ

Table 12 shows these values of M for the magma densities that wouldnot otherwise be erupted. Note that these are absolute minimumvalues for the depth of the mantle melt source below the base of thecrust: these values would just raise magma to the surface but wouldnot provide a pressure gradient to allow an eruption at a finitedischarge rate. We can readily find the depth below the base of thecrust, M′, at which melt must be generated to provide any chosenpressure gradient dP/dz:

MV= C dP=dz − g ρcm − ρð Þ½ �= g ρm − ρð Þ− dP=dz½ �f g ð17Þ

Table 12 shows the values ofM′ needed to produce dP/dz=500 Pam−1,typical of the values found earlier; they are of order tens of kilometers.Deciding which of these depths might be plausible involves considera-tions of both thermal and petrological modeling of the martian mantlewhich are beyond the scope of this paper. It seems clear, however, thatwhateverweassumeabout the locationofmagmareservoirs feeding thesurface eruptions, it will be reasonable (with a factor of order 2) toassume a value of dP/dz=500 Pa m−1.

6.4. Subsurface dike geometry and magma rise speed

We can now relate the typical pressure gradient driving magmamotion deduced above, ~500 Pa m−1, to the geometry of the dikefeeding any one segment of the erupting fissure. The analysis isconveniently done in terms of the magma volume flux per unit lengthalong strike of the feeder dike, denoted F. This is obtainedbydividing thevolume flux of a typical flow by the active length, X, of the fissure thatproduced it. For the Central vent complex, where the relationshipsbetween theflowunits and the vent system is particularly clear (Fig. 6b),the typical lengths along strike of fissure segments feeding flows are~930 m. With the typical volume fluxes found above, ~5000 to10,000 m3 s−1, this implies values of F of order 5 to 10 m3 s−1 m−1. Ifthe subsurface dike has width wd, and the magma is flowing up it atmean speedud along the active horizontal lengthX, then clearly the totalvolume flux is equal to (ud wd X) and so F is equal to (ud wd X)/X, i.e.:

F = udwd ð18Þ

ud and wd are related by analogs of Eqs. (2a) and (2b):

ud = w2d dP=dz

� �= 12 ηð Þ if the flowmotion is similar ð19aÞ

ud = wd dP=dzð Þ= f ρð Þ½ �1=2 if the flowmotion is turbulent ð19bÞ

As is the case for surface flows, whichever of these equations gives thesmaller value of ud is the correct one to use, a result that can bechecked by evaluating the Reynolds number, in this case:

Re = 2ud wd ρð Þ= η ð20Þ

The friction factor needed in Eq. (19b) is again given by Eq. (4) butwith Eq. (20) defining Re.

We begin by noting that the magma flow in the dike system mustby near the transition from laminar to turbulent flow. If Eq. (18) isused to substitute for (ud wd) in Eq. (20) we have:

Re = 2 F ρð Þ= η ð21Þ

Using the earlier findings that η is ~100 Pa s and that F lies between 5and 10m3 s−1m−1, thenwith ρ=~2500 kgm−3 we see that Re=250when F=5 m3 s−1 m−1 and Re=500 when F=10 m3 s−1 m−1. But

the change from laminar to turbulent flow occurs at Ret=316, and sothe range found for F spans the transition.

Now combining Eqs. (18) and (19a) we find that if the motion islaminar:

ud = F2dP=dz� �

= 12 ηð Þh i1=3 ð22aÞ

wd = 12η Fð Þ=dP=dz½ �1=3 ð22bÞCombining Eqs. (18), (19b), (21) and (4) when themotion is turbulentwe have, after a little algebra:

ud = 1:5489 dP=dzð Þ1=3 F5=12h i

= ρ1=4 η1=12� �

ð23aÞ

wd = 0:6456 ρ1=4 η1=12 F7=12� �

= dP=dzð Þ1=3 ð23bÞ

As before we use ρ=~2500 kg m−3 and η=100 Pa s. Then, whenF=5 m3 s−1 m−1, the laminar formulae yield ud=2.18 m s−1 andwd=2.29 m, and the turbulent formulae yield ud=2.32 m s−1 andwd=2.16 m; the appropriate values are the laminar solution,ud=2.18m s−1 andwd=2.29m.When F=10m3 s−1m−1 the laminarformulae yield ud=3.47 m s−1 and wd=2.88 m, and the turbulentformulae yield ud=3.09m s−1 andwd=3.23m; the appropriate valuesin this case are the turbulent solution, ud=3.09m s−1 andwd=3.23m.In summary, unless the magma properties and driving pressureconditions are very different from those derived earlier, these eruptionswere fed fromdepthbymagma rising at 2–3ms−1 through dikes 2–3mwide.

7. Eruption conditions at the surface fissure

Finally we turn to the implications of our earlier assertion that thepresence of ramparts along the sides of much of the length of theCentral vent complex fissure (see Fig. 7) implies that these eruptionswere explosive, with flows fed by the accumulation of hot spatter fromvigorous fire fountains. In general, lava fountains have an opticallydense core and a translucent outer envelope. Clasts in the opaque corecannot radiate to the surrounding sky— they simply exchange radiantheat with one another. A certain amount of atmospheric gas isentrained into the fountain and passes up into a convecting plumeabove it, carrying away those clasts small enough to have low terminalvelocities in the gas. This gas stream removes some heat from all of theclasts within the fountain, so the core is not quite isothermal.Nonetheless, the larger clasts travelling and landing within the coremay be able to coalesce on reaching the ground to generate a rootlesslava flow, and we infer that this is how the flows studied here wereformed. In the translucent outer envelope, clasts both radiate heat tothe surroundings and lose heat to the entrained gas, and hence arecooler on landing. These clasts may be able to partially or completelyweld to one another, but cannot deform to the point of flowing; theyare candidates for forming spatter ramparts.

A model of lava fountain opacity was developed for axi-symmetricpoint-source vents by Wilson and Head (1981) and Wilson and Keil(1996) and this was extended to a linear fissure vent geometry byWilson and Head (2001). The model shows that the thickness, Λ, ofthe outermost, translucent envelope, which we equate to the width ofthe spatter ramparts, is given by:

Λ = 2π/RUð Þ= 3 Fð Þ ð24Þ

where ϕ is the median grain size of the ejected clasts, R is theirmaximum range in the opaque part of the fountain, equated here tothe maximum distance from the fissure at which clast coalescence toform flows is interpreted to have taken place, U is the average speed ofthese clasts as they leave the vent, and, as before, F is the volume flux


per unit length along strike of the fissure. We can measure Λ and R:Fig. 8a shows that Λ is mainly in the range 100 to 150 m and that R istypically 1000 to 1500 m. However, we now need to relate thesemeasurable quantities to ϕ and U, and to relate all of these quantitiesto our value of F estimated from the surface lava flow lengths. Thisrequires consideration of the likely volatile content of the eruptingmagma, because it is the exsolution of volatiles, and the expansion ofthe resulting gas bubbles to the point where the magma is disruptedinto a spray of gas and pyroclastic droplets, that determines themedian droplet size ϕ. Similarly the expansion of the released gasdetermines the ejection speed, U, of the droplets and hence their finalrange R.

A one-dimensional model of the consequences of volatile releasefrom magmas is given by Wilson et al. (1980) and Wilson and Head(1981). It utilizes the approximation that the pressure in the magmapassing though the system, both before and after its disruption intopyroclasts, is equal to the lithostatic pressure in the host rocks; inother words there is nowhere any net stress across the walls of thedike and vent system. This leads to entirely plausible dike and ventgeometries for eruptions on the Earth. However, when it is applied tomartian conditions (Wilson and Head, 1994; Mitchell, 2005), it isfound that, for almost all realistic combinations of exsolved magmagas content and erupted mass flux, the mixture of pyroclasts and gasleaving the vent cannot expand enough to reach atmosphericpressure, but instead emerges as an over-pressured, choked flow inwhich the exit speed is constrained to be equal to the speed of soundat the exit pressure and temperature. Although some aspects of thebehavior of such systems have been addressed by Kieffer (1984), thesubsequent decompression of the jet to atmospheric pressure througha series of shocks has not so far been treated in detail for volcanicsystems (Mitchell, 2005).

Here we approximate the vent conditions using a modification ofthe computer code described in Wilson and Head (1981) to computethe upward acceleration of the rising gas–pyroclast mixture producedby magma disruption. The pressure in the mixture is set equal to thelithostatic pressure in the host rocks until the depth where the rapidlyincreasing average speed of the mixture becomes equal to the moreslowly decreasing local speed of sound. A limit is then placed on therate at which the edge of the region occupied by the mixture canexpand, that limit being the angle whose tangent is equal to thereciprocal of the local Mach number (Kieffer, 1984; Mitchell, 2005). Itis tacitly assumed that the walls of the fissure can adjust to allow thisexpansion to take place. This should not be a problem in the presentcase where the uppermost part of the magma pathway consists of theaccumulating deposits from earlier phases of the eruption, andretrospective comparison of the results of the calculations with theshape of the vent structure obtained from the images confirms thatthis is the case. The pressure and average speed of the gas–pyroclastmixture are tracked until the local atmospheric pressure is reached,and values are then noted for the height, H, at which equilibriumwiththe atmosphere is reached, the average kinetic energy per unit mass,K, of the eruption products, the width, W, of the jet, and the angle, θ,which its outer edge makes with the vertical.

Table 13Characteristics of explosive eruptions. The variations with exsolved magma volatile content,pyroclasts,ϕ, the terminal velocity of these pyroclasts in the gas,Ut, thewidth of the volcanic jetW, the height at which this occurs, H, the angle from the vertical made by the edge of the jet,component of typical ejected pyroclasts, Uh, the total speed of these typical clasts,U, the predictwidth of the spatter ramparts, Λp5 and Λp10, for two values of the erupted volume flux per unit

N Ug ϕ Ut W H θ(mass%) (m s−1) (mm) (m s−1) (m) (m) (°)

0.50 194 3.5 46 327 468 17.80.25 129 2.8 31 252 337 18.80.10 72 2.2 18 176 215 21.1

The energy per unit mass K is then partitioned amongst all of theeruption products, i.e. the gas phase together with a range of pyroclastgrain sizes, to determine the velocities of individual clast sizes. Themethod, described byWilson (1999), utilizes the fact that the speed ofeach clast size is less than the gas speed, Ug, by the terminal velocity ofthat clast size in the hot volcanic gas. The terminal velocity is afunction of the gas pressure (by now that of Mars' atmosphere,~500 Pa) and gas temperature (still close to magmatic, say 1450 K).The size distribution adopted for the pyroclasts ranges from thelargest clast that can just be transported out of the vent to a size twoorders of magnitude smaller, and to ensure computational accuracy 21size classes are defined in this range. Unfortunately, althoughpyroclasts have been detected at one location on Mars (Squyreset al., 2007), there is not enough information to obtain a representa-tive size distribution. We have therefore assumed that equal masses ofmaterial reside in each size interval. This is a simpler approximation tothe mass distribution than that adopted to model pyroclast dispersalin the martian atmosphere by Wilson and Head (2007) but isadequate for the present purpose.We take the typical ejection velocityof clasts to be that of clasts whose diameter, ϕ, is half the diameter ofthe largest clast that can be supported in the gas stream, on the basisof analysis of ejection of pyroclasts in well-documented terrestrialexplosive eruptions (Wilson, 1999). Subtracting the terminal velocityof that clast size, Ut, from the upward component of the gas velocitygives the vertical velocity Uv, and multiplying the total velocity of thegas stream by the sine of the angle between the edge of the jet and thevertical gives the horizontal velocity component Uh of typical clastsleaving the edge of the jet. Given that the clasts are projected from apoint at height H above, and horizontal distanceW/2 from, the axis ofthe fissure vent, and that they land close enough to the vent that thecurvature of the planetary surface can be neglected, a simple ballisticanalysis shows that they land at a predicted range Rp given by:

Rp = W = 2 + Uh Uvð Þ= g + Uh = gð Þ U2v + 2 g h

� �1=2 ð25Þ

Eq. (24) can then be used to give the predicted value of the opaqueenvelope thickness of the lava fountain, and hence the spatter rampartwidth, Λp, as:

Λp = 2π/Rp U2v + U2

h

� �1=2� �

= 3 Fð Þ ð26Þ

The computation is carried out by choosing themagmavolume fluxper unit length along strike of the dike system, F, the magmatic liquidproperties (density 2500 kg m−3, viscosity 100 Pa s and temperature1450 K, though the exact values used have a very small effect on theresults), and, critically, the mass fraction of volatiles in the magma, n.Based on analyses of martian meteorites (McSween et al., 2001), weassume that water is the commonest volatile in mafic martian rocksand use this as a proxy for the total volatile content; a wide range ofvalues of n was explored. Table 13 shows some results for ourestimated value of F=5 to 10 m3 s−1 m−1. For each assumed magma

n, of the speed of the erupted magmatic gas stream, Ug, the typical diameter of ejectedleaving the vent at the heightwhere its pressure becomes equal to that of the atmosphere,θ, the vertical velocity component of typical ejected pyroclasts, Uv, the horizontal velocityed values of the distance out to the edge of the zonewhere pyroclasts land hot, Rp and thelength along strike of the fissure, F=5 and 10 m3 s-1 m-1.

Uv Uh U Rp Λp5 Λp10

(m s−1) (m s−1) (m s−1) (m) (m) (m)

149 59 160 5100 1170 58598 41 106 2440 295 14854 26 60 925 49 25

Fig. 11. The predicted variation with magma volatile content, n, of (a, left) the maximum range, Rp, to which pyroclasts in the inner parts of fire fountains can be thrown while stilllanding hot enough to coalesce into lava flows and (b, right) thewidth, Λp, of the spatter ramparts formed by pyroclasts in the outer parts of the fire fountains which land cool enoughto avoid mobilization. Curves in (b) labeled Λp5 and Λp10 correspond to erupted magma volume fluxes per unit length along strike of the fissure of 5 and 10 m3 s−1 m−1, respectively.


volatile content n we give values for Ug, ϕ, Ut, W, H, θ, Uv, Uh and U asdefined above, and for the key predictions: the distance out to the edgeof the zone where pyroclasts land hot, Rp, and the width of the spatterramparts, Λp. We now compare these predicted values with thecorresponding observations. Fig. 11a and b show how Rp and Λp varywith n, the curves being labeled Λp5 when F=5 m3 s−1 m−1 and Λp10

when F=10 m3 s−1 m−1. The observed range of R, 1000 to 1500 m,corresponds to n in the range 0.11 to 0.16 mass %. If F=5m3 s−1 m−1,the observed range of Λ, 100 to 150 m, corresponds to n in the range0.14 to 0.18 mass % and if F=10 m3 s−1 m−1 the observed range of Λcorresponds to n in the range 0.21 to 0.27mass %. The range of values ofn implied by the observed values of R is very similar to the range ofvalues of n implied by the observed values of Λ if F=5m3 s−1m−1.Wetherefore conclude that Fwas probably closer to 5 than 10m3 s−1 m−1

in these eruptive episodes, and that the equivalent magma watercontent nwas probably near the middle of the range 0.1–0.2 mass %.

8. Discussion

Fissure eruptions of the type identified here are rare on Mars. Theeruptions we have studied appear to post-date the larger young lavaflows that comprise the surface units within central Tharsis (Scott andTanaka,1980), and all of the lava flow units that we havemeasured aremuch thinner (by as much as an order of magnitude) than is typicallyfound either within Tharsis or the other main volcanic region onMars,Elysium Planitia. The eruptions may, however, offer a useful insightinto the earliest phase of activity at some of the smaller volcanicconstructs on Mars because the total volume of lava erupted at thisfissure is relatively small compared with that at the main constructsand so the original structure of the vent has not been buried. Due to itsclose proximity and its apparent alignment with the vents describedhere (Fig. 2a), it is relevant to ask if any of the three eruptive centers iscomparable in activity to Jovis Tholus. This volcano lies only 45 kmfrom the summit of the westernmost of our vents, and is in almoststraight alignment with all of these vents. Plescia (1994) interpretedJovis Tholus to be a shield volcano of similar origin to two other smallvolcanoes within Tharsis, namely Ulysses and Biblis Paterae. JovisTholus is ~860 m high and is the least complex of these three smallshield volcanoes, with simple repeated outpouring of lavas andcaldera-forming events. The flanks of Jovis Tholus show none of thevarious lava flows types that the vents to the east possess (i.e., nospatter deposits, channel-fed flows, or discrete lobate flows) andappear to be comprised of undifferentiated flows that are much moreheavily mantled by dust (and so Jovis Tholus is probably significantlyolder than the other vents discussed here). However, on the north-

eastern flank of Jovis Tholus, a fissure has been recently active, and aprominent lava fan extends to the north for ~26 km. The small fissurevent constructs that we have identified are different in size from thesmall shields identified on the Tharsis Montes by Mouginis-Mark andChristensen (2005) and by Bleacher et al. (2007) because of thegreater volume of individual flows, as well as the linear nature of someof the vent systems. We can, however, envision a spectrum of sizes ofconstructs, with the small shields on the Tharsis Montes grading intothe western and eastern vent complexes, with a further increase involume producing a volcano comparable to Jovis Tholus.

The rheological properties deduced for the flows from these fissurevents are not typical for Mars. We have commented above that theflows are unusually thin. The thickness of a flow is not just a functionof its rheology; as shown by Eq. (1), the topographic slope onwhich itis emplaced is also involved. However, taking account of this, we findyield strengths deduced from flowmargin thicknesses of ~100 Pa, oneto two orders of magnitude smaller than values deduced for flowselsewhere onMars. Similarly we have found lava viscosities of ~100 Pas, again at least an order of magnitude less than found for flowselsewhere. Taken together these results imply some combination ofunusual composition, unusual pre-eruptive storage, or unusualeruption conditions. The presence of spatter ramparts along parts ofthe fissure of the Central vent complex, together with the streamlinedstructure of the levees of the proximal parts of many of the flow units,have been used to infer that these flows were commonly fed by firefountain activity driven by volatile release. Such activity haspreviously been proposed for Mars (Wilson and Head, 1994) butuntil this case has not been confidently identified. Although thecentral parts of optically thick fire fountains do not lose heat, therootless lava flows formed from the hot clasts descending from themare inevitably contaminated with cooler clasts descending from theouter, transparent parts of the fountains, so it cannot be argued thatthese flows retained more heat than usual on being erupted — ifanything they must have lost more heat than if they had been fed byquiet overflow of volatile-poor magma from the fissure. Thus unusualeruption conditions do not seem an attractive explanation for the lavarheology.

The possibility of unusual magma storage prior to eruption largelyhinges on whether the magma was being erupted from a shallowmagma reservoir or directly from a source region in the mantle, or atleast at the base of the crust. Clearly, rapid eruption either directlyfrom amantle source region or from a reservoir at the base of the crustis likely to lead to hottermagma than eruption from a crustal reservoirin which significant cooling (and consequent chemical evolution) ofmagma can occur prior to eruption. Unfortunately, we have seen that


there seems to be nothing that can be deduced from the analysis of theeruption dynamics alone that distinguishes between shallow anddeep reservoirs. However, the low viscosity that we have found to beconsistent with the flowmorphologies may itself be the best evidencethat these lavas are more likely to be primitive mantle melts erupteddirectly from great depth than the products of prolonged shallowcrustal storage.

This uncertainty in the depth of origin of the magma means that itis not possible to infer the overall eruption rate of the vent complexes.We do not, for instance, have any stratigraphic information thatconfirms or refutes the idea that more than a small (b10 km) segmentof the entire fissure system was active at any instant in time. Nor canwe determine if the vent complexes were active simultaneously orerupted thousands (millions?) of years apart. Had we been able toinfer a chronology to the activity at the vents, then it might be possibleto identify an underlying structural control to the distribution of theactivity.

Our review of image data from THEMIS has so far failed to identifyany comparable fissure systems that produced thin flows of limitedextent elsewhere on Mars. The northeastern plains of Tharsis containmany lava flows N100 km in length that we infer were erupted fromfissure systems that could have been comparable to the ventsdescribed here, but clearly these NE Tharsis fissures were far moreprolific both in terms of the spatial distribution and the volumeerupted from each segment of the fissure. Thus we conclude that theexamples studied here may be unique on Mars, although additionalsearches of older volcanic areas (Syrtis Major, Elysium Planitia andHesperia Planum) may also show greater diversity than is currentlyaccepted. Particularly with the advent of image data with even higherspatial resolution than the THEMIS VIS instrument (e.g., HiRISE andContext Imager data from Mars Reconnaissance Orbiter), an ongoingsearch for fissure-fed flows on Mars appears to be warranted.

9. Summary

(1) We have identified and mapped six vent systems in easternTharsis, east of the volcano Jovis Tholus. All of them have builtsmall shields rising to ~50–85 m above the level of thesurrounding topography and have fed numerous channelizedlava flows. All but one are linear fissures a few to ~20 km inlength, and the remaining one probably began life as a fissurethough its summit evolved into a lava pond that bears aremarkable similarity to the Kupaianaha lava pond at Kilaueavolcano, Hawai'i.

(2) The maximum length of individual lava flows from thesefissures is ~30 km, with typical lengths being 15–20 km. Flowwidths are typically within a factor of 2 of 1500 m.

(3) Flow thicknesses are all b5 m, which is an order of magnitudethinner than essentially all flows previously identified elsewhereonMars. Photoclinometric profiles at ~20mhorizontal scale havebeen constructed across one of the longer flows to confirm thethickness values derived from ~300 m scale MOLA data.

(4) Smooth material, interpreted to be spatter from fire fountaining,is located on the rims of some of the fissures. Some of the spatterremained in place, and some formed either short (b5 km) flowsor merged close to the fissure to form longer flows.

(5) There appears to have been some temporal evolution of theflow field, as the oldest parts of the basement at each center arebuilt from a series of compound flows that cannot besubdivided into individual flows.

(6) Modeling of the dynamics of emplacement of the flows stronglysuggests that they consist of fluid (viscosity ~100 Pa s), lowyield strength (~100 Pa) lava erupted at typical rates of~5000 m3 s−1 from ~1 km-long segments of the fissures.Flow speeds were ~1–2 m s−1 and durations of emplacementof individual flow units were a few hours.

(7) Modeling of the emplacement of spatter ramparts suggests thatthe volatile content of the magma, expressed as equivalentwater content, was ~0.1–0.2 mass%, enough to cause explosiveactivity with eruption speeds of ~70–120 m s−1 driving firefountains to 650–1900 m above the vent.

(8) Modeling of the source depth of the magma is inconclusivebecause a wide range of source depths leads to similar pressuregradients driving magma to the surface (within a factor of ~2 of~500 Pa m−1), with implied dike widths of 2–3 m and magmarise speeds, beneath the depth of magma fragmentation byexpanding gas bubbles, of ~2–3 m s−1.

(9) The eruptions described here may characterize a style ofvolcanism on Mars that can only be identified with imagespatial resolutions better than 20 m/pixel. It is thereforepossible that additional searches of other volcanic areas (SyrtisPlanum, Elysium Planitia, and Hesperia Planum) using newimage data may also show a greater diversity of eruptiveactivity than is currently accepted.

NotationSymbol Definition, valueC thickness of planetary crustF magma volume flux per unit length along strike of feeder dikeGz critical value of dimensionless Grätz number, 300H height at which erupted gas–clast mixture reaches atmo-

spheric pressureK average kinetic energy per unit mass of eruption productsL length of lava flow unitM minimum depth of melt source below base of crustM′ depth of melting below base of crust to provide given

pressure gradientP lithostatic pressure in crustQt lava volume flux from vent at boundary between laminar

and turbulent flowQc volume flux of cooling-limited lava flow unitR observed maximum range of ejected clasts in opaque part of

the fountainRp predictedmaximum range of ejected clasts in opaque part of

the fountainRe Reynolds number for lava or magma motionRet Reynolds number at boundary between laminar and

turbulent flowU average speed of erupted pyroclasts leaving the ventUg speed of magmatic gas leaving ventUh horizontal velocity of pyroclastUt terminal velocity of pyroclastUv vertical velocity of pyroclastV volume of lava flow unitW width of gas–clast jet leaving the ventX active length along strike of erupting fissured depth of lava flowdP/dz pressure gradient driving magma motionf friction factorft friction factor at boundary between laminar and turbulentflowg acceleration due to gravity, 3.72 m s−2

nCO2 solubility of CO2 in basalt as mass fractionu mean speed of lava flowud mean speed of magma flowing up dikeut lava flow speed at boundary between laminar and turbulent

flowv pore space volume fraction in crustal rocksvs pore space volume fraction in surface rocksw width of lava flow unitwd width of dikez depth below surfaceΔP crust base lithostatic pressure minus weight of magma in


dike reaching surfaceΛp predicted value of Λ, thickness of outermost, translucent

envelope of fire fountainΛp5 value of Λp when F=5 m3 s−1 m−1

Λp10 value of Λp when F=10 m3 s−1 m−1

α slope of ground beneath lava flowβ dimensionless factor depending on open channel or tube

flow regime, 4η viscosity of lavaηt lava viscosity at boundary between laminar and turbulentflowθ angle which outer edge of erupting gas–clast jet makes with

verticalκ thermal diffusivity of lava, 7×10−7 m2 s−1

λ gravity-independent compaction parameter, 1.18×10−8 Pa−1

ϕ median grain size of erupted pyroclastsρ lava density, 2500 kg m−3

ρc local density of planetary crustρcm mean density of planetary crustρcs solidified magma densityρ0 lithosphere density at depth where all pore space vanishesτ duration of emplacement of lava flow unit

Acknowledgements

This work was supported in part by NASA grant NNG05GQ54Gfrom NASA's Mars Data Analysis Program. We thank Phil Christensenand the THEMIS Science Team, especially Laurel Cherednik, KellyBender, and Abdreas Dombovari, who were responsible for thetargeting of THEMIS instrument to collect the extensive coverage ofthe study area, and to the creators of the THEMIS web site (http://themis.asu.edu) that facilitated our access to these data. Helpfulcomments by Jim Zimbelman and an anonymous reviewer were muchappreciated. This is HIGP Publication No.1716 and SOEST ContributionNo. 7714.

Appendix 1. Photoclinometric determination of local topography

The first stage of this process is to download the THEMIS visible RDRproduct from theASUTHEMISwebsite (themis.asu.edu). This is the level1 calibrated radiance product in PDS format. Using a script, we applythree isis commands: thm2isis, thmvismc, and levpt. These convert thePDS format file to an ISIS cube, project the data into a map projection(simple cylindrical, planetocentric, central meridian of zero degrees),and generate the geometry information for the sample and line spatialresolution, incidence angle, emission angle, azimuth angle, and solarazimuth that are necessary for the photoclinometry operation. Thegeometry information is written to an associated ASCII file.

We next determine the location of each raw MOLA elevationmeasurement. This is done by means of an in-house producedprogram called Molapts. The Molapts program is a Java GUI-basedprogram that displays a user specified level 2 THEMIS scene,processed with the script described above, and overlays the appro-priate MOLA shots upon the THEMIS image (see, for example, Fig. 7).The Molapts program queries a MySQL database that we have createdwhich contains the latitude, longitude, elevation and orbit number forall 600,000,000 MOLA shots. By reading the starting latitude andlongitude and the spatial resolution from the ISIS level 2 cube header,we overlay MOLA shots that are within the latitude and longitudeboundaries of the cube at their correct position. These shots arerepresented on the grayscale rendering of the image (the open circlesshown in Fig. 7). The user can click on any MOLA shot, represented ascircle outlines, and see the latitude, longitude, and elevation of thatspecific point.

Because of a timing uncertainty in the acquisition of the THEMISimages (see http://isis.astrogeology.usgs.gov/Isis2/isis-bin/themis_-processing.cgi#9), there can be a small north or south offset between

where a MOLA point is placed and where it should be placed. Byinspection of the MOLA shots over topographic features, a user canmake small north or south adjustments to the placement of the imageunder the shots to minimize the offset caused by the timinguncertainty.

The photoclinometry program is designed to work with threeknown, nearly collinear, elevations defining the profile, and with thecalibrated radiance available as THEMIS image data. The knownelevations are individual MOLA shots. The user selects the start,middle and end point MOLA shots from three different orbits, theseshots being typically aligned along the direction of illuminationwithinthe scene. With the geometry known, the fitting is a matter of findingthe topographic “flat” pixel radiance. The program uses the methoddescribed in Glaze et al. (1999) to calculate the profile elevations. Theprogram iterates through the shadow or dark radiance as well as theflat pixel radiance, i.e. there is a flat pixel loop within a dark pixel loop.The procedure consists of starting with the elevation at the startingMOLA shot, calculating the slope at each pixel along the profile fromthe radiance at that location, and then multiplying the slope by thedistance between pixels along the profile to get the change inelevation from the previous pixel. This continues along the profile,determining an elevation at each profile location. The best fit isselected as the combination of dark pixel and flat pixel radiance thatminimizes the difference between the photoclinometrically deter-mined elevation and theMOLA elevation at the three known locations.The resulting profile can be output to a file and is also displayed in awindow on the photoclinometry control window. Selection of parallelprofiles using adjacent shots within the MOLA profiles produces a setof topographic profiles that can be stacked together to show thetopography at the spatial resolution of the THEMIS VIS images (18 m/pixel), as shown in Figs. 7b, 8b and 9b.

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Fissure eruptions in Tharsis, Mars: Implications for eruption conditions and magma sources

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