slabby pahoehoe from the western deccan volcanic province: evidence for incipient pahoehoe^aa...

Slabby pahoehoe from the western Deccan Volcanic Province:evidence for incipient pahoehoeâa transitions

Raymond A. Duraiswami �, Gauri Dole, Ninad BondreDepartment of Geology, University of Pune, Pune 411 007 India

Received 1 June 2001; received in revised form 30 July 2002; accepted 30 July 2002

Abstract

The pahoehoeâa transition for a flow exposed near Bodshil village from the western part of the Deccan VolcanicProvince (DVP) is reported for the first time. The V1-km-long Bodshil flow issued as a small sheet from a pre-existing lobe. Near the source, the crust is characterised by numerous squeeze-ups. A number of gaping fractures,parallel to sub-parallel to the flow direction, are exposed on the surface in the medial portion of the flow. About 800m away, the flow completely transforms to slabby pahoehoe. The terminal portion of the flow is characterised byconcentrations of slabs, blocks and lava balls. The size and concentrations of the slabs and lava balls appear toincrease along the length of the flow. Petrographic studies reveal a dominant hypohyaline texture. The flow core iscoarse and is characterised by plagioclase set in a glassy matrix. The presence of clinopyroxene in addition toplagioclase and glass distinguishes the crust and interslab crust from the core. On the basis of mineralogy, atemperature range of 11465 15‡C to 11695 15‡C is inferred for the Bodshil flow. Increased vesicle deformation acrossthe transition is discernible and an average D-value of 6 0.4 indicates moderate strain rates during emplacement. Inlight of the morphology and petrography, the cooling history and the mode of emplacement of the Bodshil flow isdiscussed. The flow originated as a small toe at the leading edge of a pahoehoe flow, and grew into a sheet by themechanism of inflation. Continuous inflation caused the brittle crust to uplift and produce a network of inflationclefts that were subsequently occupied by squeeze-ups. Temporary stagnation of the flow due to cessation of lavasupply or storage allowed the crust to grow and thicken. Renewed movement of the stored and cooled lava to the flowfront at a fairly high volumetric rate was responsible for the initial disruption of the crust. High rates of crustaldisruption induced higher rates of degassing and cooling, which resulted in rapid crystallisation of the fluid core.Increase in crystallinity lead to the onset of yield strength, and it is envisaged that at least the terminal parts of theflow behaved as a Bingham fluid. The Bodshil flow is unique to the DVP because it is the first to record slabbypahoehoe and provide evidence for the incipient transformation of basaltic lava from pahoehoe to aa.< 2003 Elsevier Science B.V. All rights reserved.

Keywords: volcanology; in£ation; slabby pahoehoe; pahoehoeâa transition; Deccan Volcanic Province

1. Introduction

Surface morphology and structures distinguishpahoehoe lava from aa (Macdonald, 1953;Wenthworth and Macdonald, 1953). Pahoehoe

0377-0273 / 03 / $ ^ see front matter < 2003 Elsevier Science B.V. All rights reserved.doi:10.1016/S0377-0273(02)00411-0

* Corresponding author. A/6, Gurudut Housing Society,Kalewadi Phata, Srinagar, Rahatani, Pune 411 017, India.

E-mail address: raymond_d@redi¡mail.com(R.A. Duraiswami).

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Journal of Volcanology and Geothermal Research 121 (2003) 195^217

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(crust-dominated) £ows are characterised by asmooth, billowy or ropy surface, and containspherical vesicles. Bulbous toes, lobes, sheets,tubes and tumuli are common features associatedwith pahoehoe. Internal structure of pahoehoe£ows exhibit a typical three-tiered structure (Au-bele et al., 1988): an upper vesicular zone (crust)followed by a non-vesicular zone (core) and alower vesicular zone (basal zone). Aa (core-dom-inated) £ows are characterised by a fragmentaryand spinose surface with irregularly shapedvesicles. Lava channels and accompanying leveesare generally associated with aa £ows. The upperclinker layer (£ow top breccia) of aa grades into acentral massive layer, which may grade into athinner, less persistent, clinker layer at the bot-tom. Besides clinkers, the tops of aa £ows mayalso contain accretionary lava balls. Aa lavascan be distinguished into proximal-type and dis-tal-types (Rowland and Walker, 1987). Proximal-type aa occurs near the vents and contains sur¢-cial rubble of scoriaceous clinkers. Distal-type aaforms s 10-m-thick lobes and occurs in the distalparts of £ow ¢elds. Their surfaces are coveredwith rubble of non-scoriaceous lava derivedfrom uprising lava from the £ow interior.Pahoehoe and aa lava £ows not only di¡er

super¢cially in surface textures but also funda-mentally in their emplacement mechanism (Kil-burn, 1993). According to Rowland and Walker(1990), pahoehoe forms at low volumetric £owrate (6 5^10 m3/s) and aa forms at high volumet-ric £ow rate (s 5^10 m3/s). Pahoehoe £ows areinvariably emplaced by endogenous processes, i.e.through the mechanism of in£ation (Hon et al.,1994). Rapidly formed units in pahoehoe £owsare almost always preserved due to slow rates oflava in£ux, so that heat loss is negligible (Marsh,1981; Cashman et al., 1994; Keszthelyi, 1994,1995). Rapidly advancing aa £ows, on the otherhand, cause continuous disruption of the £owcrust and allow the exposed core to lose heatquickly. The rapid heat loss leads to high ratesof crystallisation (Cashman et al., 1997), increas-ing lava crystallinity and lava viscosity, both con-tributing to changes in lava rheology (Dingwelland Webb, 1989; Webb and Dingwell, 1990; Pin-kerton and Stevenson, 1992; Crisp et al., 1994).

Proximal-type aa advances with a rolling caterpil-lar-track motion at rates greater than 1 m/min onlow angle slopes, and the distal-type aa generallymoves at a slower rate of 1 m/10 min on low angleslopes and does not advance with a rolling cater-pillar-track motion (Rowland and Walker, 1987).Most basaltic lavas from Hawaiian eruptions

are initially of the pahoehoe type (Swanson,1973; Lipman and Banks, 1987), and only a fewhave transformed to aa during £owage. Althoughthe pahoehoeâa transitions have been well-docu-mented amongst recent Hawaiian £ows (Petersonand Tilling, 1980; Cashman et al., 1999; Polacciet al., 1999), descriptions of similar transitions arevirtually non-existent in the literature for ancientContinental Flood Basalt (CFB) provinces ^ espe-cially in the Deccan Volcanic Province (DVP), forwhich volcanological studies have been grosslyneglected. Previous studies suggest that the pahoe-hoeâa transition requires su⁄ciently high strainrates (Peterson and Tilling, 1980) as well as acrystallinity-induced onset of yield strength equiv-alent to the imposed stress (Cashman et al., 1999).According to Peterson and Tilling (1980), thephysical process of transition from pahoehoe toaa is accomplished by: (a) spontaneous formationof relatively sti¡ ‘clots’ in parts of the £owing lavawhere shear rates are high; (b) fragmentation andimmersion of solidifying pahoehoe crustal slabsinto the £ow interior; and (c) by sudden renewedmovements of the lava stored and cooled withinsurface reservoirs. The transformation of the Bod-shil £ow, described in the present paper, proba-bly incorporates all the above-mentioned mecha-nisms. Textural changes across pahoehoeâatransitions in the Bodshil £ow include an increasein microlite crystallinity and vesicle deformationsimilar to that reported by Polacci et al. (1999).The Bodshil £ow, therefore, quali¢es as a genuinecase of pahoehoeâa transition from an ancientCFB, and is probably the ¢rst of its kind in theDVP.

2. Flows in the Deccan Volcanic Province

The 64^65 Ma DVP occupies more than500 000 km3 area in parts of western and central

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India and is one of the larger CFB provinces inthe world. Prof. G.P.L. Walker introduced theterms ‘compound’ and ‘simple’ to basalt unitsfrom the DVP during his visit in 1969. Sincethen ‘compound £ow’ in the DVP has come tomean a £ow consisting of several to many sub-equal-sized lobes that may or may not be verti-cally overlapping (see Keszthelyi et al., 1999b),while a ‘simple £ow’ consists of a single coolingunit. According to Keszthelyi et al. (1999b), mostsimple £ows may actually be large sheet lobeswhich, when traced over a considerable distance,may terminate against other lobes. Compoundpahoehoe £ows sensu lato (Walker, 1971) are ex-posed around an elliptical region (Fig. 1) and‘simple’ £ows sensu lato are dominant in the pe-ripheral regions (Deshmukh, 1988) of the prov-ince. Vents are rarely exposed in ancient CFBprovinces (Self et al., 1997), and are virtually un-documented in the DVP. The presence of com-pound pahoehoe £ows in the Nashik^Igatpuri re-gion suggests that this was a major eruptivecentre. Further support for this region being thesource area comes from the observed oversteppingof geochemical stratigraphic formations (Cox,1983; Subbarao et al., 1988; Mitchell and Wid-dowson, 1991), the location of the maximumthickness of the lava pile and the regional struc-ture of the western DVP.In the DVP, compound pahoehoe £ows display

a variety of volcanological features (Phadke andSukhtankar, 1971; Duraiswami et al., 2001, 2002)akin to their Hawaiian counterparts. The simple£ows, on the other hand, are characterised by

vesicular crust, with or without £ow top breccia,thick, non-vesicular, dense but jointed core and asharp glassy base.Recent studies in the DVP have underlined the

importance of in£ation in the emplacement of pa-hoehoe £ows (Keszthelyi et al., 1999b; Bondre etal., 2000). Many £ows in the province are mistak-en for aa due to their blocky or bouldery appear-ance ^ an artifact probably related to weatheringand erosion of the pahoehoe crust and exposedjointed core. The spatial distribution of com-pound pahoehoe £ows and simple £ows in theDVP (Fig. 1) suggests that the compound pahoe-hoe £ows occur in regions proximal to the hot-spot/plume trace (Mitchell and Widdowson,1991), while the simple £ows occur towards thedistal ends. However, the temporal distribution ofthe £ow types across the stratigraphy (Table 1)tells a di¡erent tale. Detailed geochemical stratig-raphy of the western part of the province (Table1) divides the V1.5-km-thick volcanic pile intotwelve formations (see Bean et al., 1986; Coxand Hawkesworth, 1985; Devey and Lightfoot,1986; Subbarao et al., 1988). Lithostratigraphicclassi¢cation ascribes the entire province a Super-group status that is divisible into the Sahyadri,Satpura, Malwa and Amarkantak groups (God-bole et al., 1996; G.S.I., 1998). Flows belongingto the Sahyadri Group are exposed in the westernparts of the province and are grouped into nineformations (Table 1). The Bushe Formation(equivalent to the Karla Formation, see Godboleet al., 1996; G.S.I., 1998) is unique to the DVP,not only because it divides the stratigraphy on the

Fig. 1. Map of the DVP (adapted after Deshmukh, 1988), showing distribution of compound and simple £ows. Inset: Map ofthe area around the village Bodshil, showing the extent of the £ow described in the text.

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R.A. Duraiswami et al. / Journal of Volcanology and Geothermal Research 121 (2003) 195^217 197

basis of petrography and geochemistry, but alsobecause it separates the predominantly compoundpahoehoe £ows belonging to the lower formationsfrom the simple £ows in the upper formations. Itis not yet clear as to why there is a change in £owmorphology across the stratigraphy but such achange probably implies a temporal change inthe nature and style of the Deccan eruptionsthrough time (Duraiswami et al., 2001). Sincevolumetric £ow rate controls lava structure (Row-land and Walker, 1990), it is tempting to attributethe peculiar distribution of £ow type to changesin the rate of lava supply, especially towards thelatter half of the volcanism. However, such anover-simplistic view would have been possiblehad there been abundance of transitional £owsin the province. In such a situation, the study oftransition of pahoehoe to aa in individual £owsshould go a long way in understanding the em-placement of such £ows and may shed light onwhy such £ows did not develop on a regionalscale.

3. The Bodshil £ow

Bodshil village (18‡37P39QN; 73‡27P37QE) lies tothe southeast of Lonavala close to the crest of theWestern Ghats (Fig. 1, inset). The £ows exposedaround the village belong to the Bushe Forma-

tion, while those exposed higher up in the hillsbelong to the Poladpur Formation. As mentionedbefore, £ows belonging to the Bushe Formationare compound pahoehoe in nature, and occur asan overlapping sequence of in£ated sheet lobes(sensu, Thordarson and Self, 1998) that vary inthickness from 1 to 10 m. Numerous smaller toes(V0.5 m long, V0.10 m thick) and lobes (s 1 mlong) form as outbreaks along in£ating fronts,especially in surface irregularities of previouslyformed lobes. It is usually di⁄cult to uniquelycorrelate a lobe from one location to anotherwithin such a sequence. The Bodshil £ow exposedto the west of the village, however, is distinctivebecause of its source, manner of extrusion and thepresence of a fragmented pahoehoe crust. This£ow lobe can be traced for a distance of 900 m,up to a point on the hill slope where it is con-cealed by the overlying £ows (Fig. 2). It is prob-able that the £ow lobe was an outbreak from anephemeral vent (henceforth referred to as the‘source’). The ¢eld geometry of the £ow indicatesthat the lava seems to have issued as a sheetrather than as a narrow stream near the source.In the proximal region, the £ow lobe is char-acterised by vesicular pahoehoe with numeroussqueeze-ups (Fig. 2A). Several distinct and hori-zontal vesicular zones impart crude banding tothe crust. Unlike other sheets and thick lobes be-longing to the Bushe Formation, the core of the

Table 1Equivalence between the established geochemical- and lithostratigraphy in the western parts of the Deccan Volcanic Province

Geochemical stratigraphy Lithostratigraphy

Group Subgroup Formation Supergroup Group Subgroup Formation

D Desur MahabaleshwarE Panhala D Mahabaleshwar M4C Wai Mahabaleshwar EC Ambenali C S PurandargadA Poladpur C A Diveghat DiveghatN A H Elephanta

Lonavala Bushe N Y KarlaKandhala A Lonavala Indrayani

B D M3A Bhimashankar T R Upper RatangadS Thakurwadi R I Kalsubai M2A Kalsubai Neral A Lower RatangadL Igatpuri P M1T Jawhar Salher

VOLGEO 2539 30-1-03


Bodshil £ow is devoid of vesicle cylinders andhorizontal vesicle sheets. Towards the northernend of the £ow, 700 m from the ‘source’, a num-ber of fractures are exposed with their long axesoriented parallel to sub-parallel with the £ow di-rection. The length of these fractures range from 9to 13 m and their average width is about 1 m.Such fractures appear to be restricted to the west-ern half of the £ow (Fig. 2B). The fractures are

invariably ¢lled with brecciated fragments and arehighly silici¢ed (Fig. 3A). A second set of gappingfractures having larger dimensions (s 15 m) areexposed in the eastern half of the £ow. These aredevoid of any slabs. The fractures appear to havean en echelon arrangement, indicating an aniso-tropic stress ¢eld in the upper crust (Reches andFink, 1988). The £ow lobe shows low (12‡^21‡)dips towards the south, and is in contrast to the

Fig. 2. Map of the Bodshil £ow showing (A) details of surface morphology of part of the £ow towards the proximal end show-ing numerous lava in£ation clefts and squeeze-ups and (B) details of the terminal part of the £ow showing fractures and slabs.

VOLGEO 2539 30-1-03


proximal parts where it shows high dips of up to25‡ in the northeast direction. Towards the termi-nal part (at around 750 m), larger fractures arepresent with lengths s 60 m and widths ofV2 m.These are invariably ¢lled with slabs of varying

dimensions (Fig. 3B). Most slabs are angular inshape and predominantly vesicular (Fig. 4A).About 850 m from the source, the £ow lobeshows a complete transformation to slabby pa-hoehoe (Fig. 3C). In this section, fragments of

Fig. 3. Field photographs of the Bodshil £ow showing: (A) small crust fragments along a fracture with silica mineralisation,(B) one of the larger fractures ¢lled by angular crust fragments, (C) slabby pahoehoe at about 800 m away from source (notethe smooth crust (interslab crust) between slabs due to upwelling of degassed lava) and (D) a solitary lobe-like fragment amongstseveral vesicular crust fragments. Length of hammer handle: 0.30 m.

VOLGEO 2539 30-1-03


varying dimensions are fused together by a red-dened glassy matrix of viscous lava and ‘rock£our’, formed by the attrition of slabs; they arehere referred to as ‘interslab crust’ to denoteyounger crust that forms after disruption of theprevious surface.Large rafts or ‘lava boats’ of pahoehoe crust

characterise the terminal portion of the £ow. Arare 2.56 mU0.48-m pahoehoe lobe (Figs. 3D,4B) is also preserved horizontally (as indicatedby the pipe vesicles). The presence of a 2.7mU1.32-m raft of vesicular crust intruded by asqueeze-up (s 0.15 m) towards the toe of the£ow resembles the crust with squeeze-ups from

Fig. 3 (Continued).

VOLGEO 2539 30-1-03


near the ‘source’. The terminal portion of the £owisV7 m thick, but a general increase in the thick-ness of the £ow is abruptly discernible at distanceof 700 m away from the source. Vesicular crustfragments far outnumber debris from the core,indicating a major contribution of slabs fromthe crust. Dimensions of fragments from the crustwere measured and plotted as a function of classinterval (Fig. 5). Most fragments have dimensionsof less than 0.50 m but can reach lengths exceed-ing 2.25 m. The dimensions of the slabs dependson the shape and size of the £ow (Moore, 1987),which in turn depends on the stress acting on thelava and its yield strength. Slab-size distributionsnot only re£ect stress conditions during £ow butmay also indicate emplacement rates (Anderson etal., 1998). Rapid fracturing of the £ow surface in

response to high extrusion rates produces smallblocks. On the contrary, low extrusion rates allowlarge slabs to form due to limited fragmentationof the crust (Anderson et al., 1998).The terminal portion of the Bodshil £ow is

also characterised by the presence of lava balls(Stearns, 1926). Larger lava balls are preservedalong the large fractures. Their diameters rangebetween 0.20 m and 0.95 m (Table 2). Smalllava balls are spherical, but larger specimens areinvariably ellipsoidal. The surface morphology ofmany lava balls is characterised by numerouscracks that ramify the surface (Figs. 6A and 7).Lava balls are devoid of any xenolithic core. Mostlava balls have a uniform crystallinity, texture andcomposition. However, some lava balls contain acoarse vesicular core and a ¢ne-grained ‘rind’ thatis less vesicular (Fig. 6B). The core and rind, how-ever, appear to be derived from the same magma.Although the morphology of the lava balls resem-bles those of bombs, they do not owe their for-mation to projection (Stearns, 1926).

4. Petrography

The petrography of the vesicular crust, inter-slab crust and core all exhibit predominantly hy-pohyaline textures. Modal analyses of individualmineral phases were determined and are presentedin Table 3.

4.1. Core

Core is coarse grained as compared to the crustand is characterised by plagioclase set in a glassymatrix (Fig. 8A). Plagioclase crystals occur aselongated laths without distinct lamellae andshow mild strain e¡ects in the form of fracturingand strain shadows. Some of the plagioclases areswallow tailed. Rarely, the microphenocrysts ofplagioclase are zoned, highly fractured and showcorroded grain boundaries. Wherever plagioclasemicrophenocrysts accumulate into aggregates,they de¢ne a glomeroporphyritic texture (Fig.8B). Although intersertal texture is predominant,a crude £ow texture is also discernible at places(Fig. 8C) and, in one of the thin sections, comb

Fig. 4. Field sketches of slabby pahoehoe from the Bodshil£ow.

VOLGEO 2539 30-1-03


layering is also observed (Fig. 8D). Dusty grainsof opaque are closely associated with glass. Dik-tytaxitic voids of various sizes and shapes arecommon. Given the absence of clinopyroxeneand the presence of plagioclase, the core temper-ature is inferred to have been between 1146515‡C and 11695 15‡C (Neilson and Dungan,1983).

4.2. Crust

The petrography of the crust is very similar tothe core, except for the predominance of zeolite¢lled vesicles in the crust. In addition to a highmodal abundance of plagioclase (Table 3), a fewcrystals of clinopyroxene are also present in someof the thin sections. However, the small size of theclinopyroxenes precludes their optical character-isation.

4.3. Interslab crust

Besides plagioclase and glass, the presenceof clinopyroxene in higher modal abundance(Table 3) distinguishes the mineralogy of the in-terslab crust from the normal crust and core.Moreover, the interslab crust is ¢ne-grained ascompared to the normal crust. Plagioclase micro-

phenocrysts are larger (s 30 Wm), highly frac-tured, and produce the glomeroporphyritic tex-ture typical of the interslab crust. Groundmassplagioclase crystals are small (6 10 Wm) and arecomparatively strain free. Clinopyroxenes gener-ally occur as small crystals (6 5 Wm) within thegroundmass (Fig. 8E). Stray crystals (s 7 Wm) areoccasionally seen in ophitic^subophitic relation-ship with groundmass plagioclase. Glass is in var-ious stages of devitri¢cation. Spherical vesicles¢lled with secondary zeolite minerals are commonto the interslab crust. A high volume percentageof glass in the interslab samples (Table 3) mayindicate low rates of crystallisation aided by rapidcooling. Based on the presence of clinopyroxenein the groundmass, a minimum temperature of11465 15‡C is inferred for the interslab crust ^ a¢gure comparatively lower than that for thecore.In the Bodshil £ow, plagioclase crystallinity

varies from 9.88 to 44.42 volume % and exhibitsan inverse relationship with glass (Fig. 9A). In-creased crystallinity greatly decreases £uidity oflava (Shaw, 1969). The initial tearing of the £owsurface results when microlite crystallinity is su⁄-cient to inhibit continuous £ow in response toshear (Cashman et al., 1999). Greater degree ofcrystallisation accompanied by decrease in £uidity

Fig. 5. Histogram depicting dimensions of crustal fragments constituting slabby pahoehoe in the Bodshil £ow.

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and high strain rates are responsible for pahoe-hoe^aa transition (Swanson, 1973; Cashman etal., 1999).

5. Vesicle deformation measurements

Initially formed spherical vesicles within lavassigni¢cantly deform because of a velocity gradientwithin £ows. Vesicle deformation, therefore, pro-vides information on strain rates related to em-placement £ow dynamics (Stein and Spera, 1992;Polacci and Papale, 1997). The vesicle deforma-tion parameter D or the ratio of vesicle length (l)to width (b) i.e. D = (l3b)/(l+b), (Taylor, 1934)was calculated for 11 samples (Table 4). In thecase of some pahoehoe samples from the Bodshil£ow, D ranges from 0.0983 to 0.2126. Vesiclesfrom interslab crust have D values between 0.0486and 0.1833. In general, D ranges between 0.1800to 0.2006 for the transitional crust. However, ahigher value of D V0.5713 is also recorded forone sample representing the transitional crust. Anapparent increase in the amount of vesicle defor-mation is therefore discernible from pahoehoe toslabby pahoehoe in the Bodshil £ow and is ad-equately re£ected in the shapes of vesicles (Fig.10A^C). Increased vesicle deformation, as seenin the present case, is a result of increased strainrate. In case of the pahoehoe samples from theBodshil £ow, the high average D value (s 0.20)can be attributed to the coalescence of vesicles(Fig. 10A). However, higher D values in thecase of the transitional samples due to coalescenceare most unlikely, as most of the vesicles arealigned in a particular direction (Fig. 10C). Thissuggests that a unidirectional simple shear wasacting on the crust of the £ow during its move-ment and led to the stretching and alignment ofthe vesicles (Polacci et al., 1999). Increased vesicledeformation, especially in the transitional sam-ples, may also be attributed to signi¢cant gasloss through a permeable network aided by bub-ble collapse (Polacci et al., 1999; Saar and Man-ga, 1999). The interslab crust, on the other hand,contains spherical vesicles (Fig. 10B) and compa-ratively lower deformation values, indicating thatthey evolved in an environment where strain rates

Table 2Dimensions and morphological descriptions of some lavaballs from the Bodshil £ow

Sr.no.

Sampleno.

Dimensions (cm) Morphological characteristics

Dmin Dint Dmax

1 BB1 15.9 17.5 18.5 Dense, non-vesicular lava balldevoid of any cracks.

2 BB2 22.5 23.9 25.8 Lava ball with small irregularvesicles. No prominent cracks.

3 BB3 32.4 34.0 36.1 Moderate-sized lava ball withfew vesicles. Prominent crackspresent along the periphery ofthe lava ball.

4 BB4 37.1 38.2 41.1 Lava ball with conchoidalsurfaces, indicative of spalling,single large crack and avesicular surface.

5 BB5 36.0 39.7 41.6 Vesicles and crackscharacterise surface of lavaball. Vesicles are ¢lled by¢brous zeolite. The cracks arenumerous and are distributedover the entire surface.

6 BB6 58.0 66.0 73.0 Large lava ball withnumerous ramifying cracksand large vesicles.

7 BB7 46.0 53.0 53.5 Lava ball with 1-cm-thickglassy rind, showingnumerous cracks and vesicles¢lled with ¢brous zeolite.

8 BB8 15.5 17.2 18.0 Small lava ball with ahomogeneous ¢ne-grainedsurface devoid of vesicles.

9 BB9 31.0 32.0 36.2 Medium-sized lava ball,ellipsoidal in shape. Surfacerough, spinose, with smallvesicles.

10 BB10 71.2 83.0 99.0 Large lava ball with smoothvesicular surface.

11 BB11 56.0 68.0 78.0 Lava ball with smooth surfaceand large vesicles.


13 BB13 40.0 56.5 57.0 Ellipsoidal lava ball withlarge, zeolite-¢lled vesicles.



16 BB16 48.5 56.0 58.0 Lava ball with a vesicularsurface and ramifying cracks.

17 BB17 60.0 63.8 68.0 Large lava ball with large(s 1 cm) elongated vesicles.

18 BB18 30.5 32.4 36.5 Lava ball with spinose surfaceand numerous cracks.

VOLGEO 2539 30-1-03


were minimal, i.e. in earlier formed fractureswhere the lava rose passively. The vesicle defor-mation parameter exhibits a positive co-relationwith crystallinity (Fig. 9B), indicating that in-creasing crystallinity a¡ects vesicle deformationand may be an important factor in pahoehoe^aatransition. The average D value of 6 0.4 is re-

corded for transitional samples from the Bodshil£ow and indicates a moderate strain rate dur-ing emplacement. Since signi¢cant degree of re-rounding and relaxation of vesicles prior to pres-ervation provides a minimum value of strain rate,it is possible that the Bodshil £ow was subjectedto a higher strain rate.

Fig. 6. Lava balls. A: Moderate-sized lava ball showing a vesicular and cracked surface. B: Section of a smaller lava ball show-ing internal distribution of vesicles. Note the coarse-grained core in contrast to the ¢ne-grained rind.

VOLGEO 2539 30-1-03


6. Geochemistry

Three samples of basalt were analysed, oneeach from the core (BOD co), crust (BOD cr)and interslab crust (BOD sq). Samples from thecrust and interslab crust were carefully chosen soas to avoid silica-¢lled vesicles. The samples werereduced to powder by using an agate mortar.Trace-element analyses for the samples were mea-sured by ICP-MS at the geochemical laboratory

of the Geological Survey of Canada. The analysesare presented in Table 5.The basalt samples from the Bodshil £ow have

high SiO2 (51.80 to 56.10 wt%, average 53.43wt%) and K2O (0.80 to 0.89 wt%, average 0.84wt%), low TiO2 (1.00 to 1.25 wt%, average1.11 wt%), Fe2O3 (10.70 to 12.30 wt%, average.11.50 wt%) and P2O5 (V0.09 wt%). MgO contentranges from 4.58 to 5.22 wt% (average 4.97 wt%).The trace element concentrations are character-ised by high Ba (228^245 ppm, average 233ppm), low Sr (150^191 ppm, average 167 ppm)and low Zr (100^118 ppm, average 108 ppm).High concentrations of large ion lithophile (LIL)elements (e.g. K, Ba, etc.), low strontium and ironconcentrations and very small concentrations ofhigh ¢eld strength (HFS) incompatible elementslike Ti, P, Zr and Y are characteristic of theBushe Formation (Bean et al., 1986). The primi-tive normalised trace-element concentrations (Fig.11) are also similar to that of the Bushe Forma-tion. Thus, the samples from the Bodshil £owshow geochemical a⁄nity to the Bushe Forma-tion. In addition, the petrography, i.e. presenceof sparse plagioclase microphenocryst in an inter-sertal groundmass and relatively low concentra-tions of opaques in the basalt, support this con-tention.

Fig. 7. Sketches of lava balls.

Table 3Volume percentage of minerals and glass from the Bodshil £ow

Sr. no. Sample no. Description Plagioclase Clinopyroxene Opaque Glass

1 BOD1c core 44.42 0.00 4.24 51.342 BOD2 crust 42.81 0.00 0.00 57.183 BOD2a crust 35.40 2.06 0.00 62.524 BOD2c core 37.92 0.00 3.41 58.675 BOD3 interslab crust 11.70 4.05 0.00 84.246 BOD3 interslab crust 14.03 1.94 0.00 84.037 BOD3d interslab crust 10.74 2.19 0.00 87.078 BOD4 interslab crust 28.18 3.48 0.00 68.359 BOD4 interslab crust 9.88 2.05 0.00 88.0710 BOD4a interslab crust 20.36 1.21 0.00 78.4311 BOD5c core 37.65 0.00 6.68 55.6812 BOD6c core 38.67 0.00 4.01 57.3213 BOD7 core 25.34 0.00 1.44 73.2214 BOD9 core 32.03 0.00 6.05 61.9215 BOD5 core 29.30 0.00 1.00 70.3916 BOD1 core 39.09 0.00 1.00 60.9017 BOD2 core 31.04 0.00 0.15 68.8518 BOD6 core 28.51 0.00 0.49 71.00

VOLGEO 2539 30-1-03


7. Mode of emplacement

The morphology and petrography of the Bod-shil £ow yields important insights into the modeof emplacement and the cooling history of an an-cient CFB £ow. The £ow probably originated as

small toes at the leading edge of a pahoehoe £ow(Fig. 12A). The toes grew to form lobes by themechanism of in£ation, i.e. the injection of mol-ten lava within a solidifying visco-elastic crust(Hon et al., 1994) and, with continuous in£ation,the lobes eventually coalesced and grew into a

Fig. 8. Photomicrographs of basalt from the Bodshil £ow. A: Randomly oriented plagioclase crystals. B: Glomeroporphyriticaggregates of plagioclase microphenocryst. C: Crude £ow texture. D: Comb-layering within the core. E: Plagioclase and clino-pyroxene in sub-ophitic association in the interslab crust.

VOLGEO 2539 30-1-03


sheet (Fig. 12B). The presence of sheet-like geom-etry of the £ow near the ‘source’ indicates em-placement over smooth pre-£ow surfaces thathad shallow slopes (Hon et al., 1994; Self et al.,1998). Continuous in£ation caused the brittlecrust to uplift, thereby producing a network oflarge and small cracks on the £ow surface, re-

ferred to as ‘in£ation clefts’ by Walker (1991).Most of these clefts were later occupied bysqueeze-ups of various sizes and shapes (Fig.2A). Subsequently, cessation of lava supply al-lowed the crust to grow and to ‘freeze’ thesheet-like morphology of the £ow, especiallynear the ‘source’. The duration of active in£ation

Fig. 8 (Continued).

VOLGEO 2539 30-1-03


of an ancient sheet £ow can be estimated from thethickness of its vesicular upper crust (Self et al.,1998). Using the empirical cooling model of Honet al. (1994), the time (t) required to grow theupper crust of a particular thickness (H) is givenby the equation t=164.8 H2. Thus, consideringthe mean thickness of the crust of the Bodshil£ow to beV1 m, it required at leastV165 hoursor V7 days to grow. The stagnation of the £owwas, however, temporary and short-lived, as isindicated by the absence of vesicle cylinders andhorizontal vesicle sheets in the core of the £ow.Oscillations in lava supply to the £ow front,

perhaps due to a £ow storing locally and/or re-treating of lava (Hon et al., 1994), are known tooccur and such sudden changes may be responsi-ble for the initial disruption of pahoehoe crust(Fig. 12C) in the Bodshil £ow (Keszthelyi et al.,1999a). Alternatively, Anderson et al. (1999) sug-gest that you can have a constant overall supplyrate, yet feed di¡erent parts of the in£ating £owat di¡erent rate or times, i.e. pulsed in£ation.Thus, pulsed in£ation with steady-supply ratecould have the same e¡ect as oscillating supplyrate on the disrupting crust. Irrespective of themechanism, high rates of crustal disruption inthe case of the Bodshil £ow allowed greater ther-

mal radiation, which in turn induced higher ratesof cooling (Cashman et al., 1997, 1999), whichaided rapid crystallisation of the still £uid core.Crustal disruption may have also aided in volatileloss (especially of CO2 and S, as these have lowersolubility in basaltic melts) and, thereby, resultedin the slight under-cooling, inferred petrographi-cally by the presence of swallow-tailed plagioclase(see Fig. 8A). Degassing can have profound ef-fects on rate of crystal nucleation and growth:crystals in basaltic melts are known to grow tosizes of 1 mm within periods of as little as a fewhours (Kirkpatrick, 1976, 1977). The crystallinityincrease leads to the onset of a yield strength andhence a transition to a Bingham £uid. An increasein viscosity and yield strength changes the stressand strain-rate conditions at which failure occurs.The ‘threshold viscosity for fracture’ depends onthe stress being applied (Kilburn, personal com-munication). Assuming that initially the stress wasfairly constant in the case of the Bodshil £ow, theincrease in viscosity led to conditions favouringfracture and ultimately to the formation of slab-by pahoehoe. Pahoehoe has a rough, transitionalsurface when crystallinities are moderate and yieldstrengths are su⁄ciently large to prevent relaxa-tion of small-scale surface roughness over short

Fig. 8 (Continued).

VOLGEO 2539 30-1-03


chilling times (Cashman et al., 1999). However,since the crust of the Bodshil £ow has not dis-rupted uniformly, we infer that the lava in thedi¡erent parts of the £ow had a considerablerange of viscosity. The brittle crust was thusundergoing widely variable rates of shear, as isamply evident from the shapes of the vesicles

and the geometry of the fractures. At about thisstage, lava ‘clots’ began to appear in the viscousparts of the £ow (see Kilburn, 1990), especiallynear some of the larger fractures. They accretedto a larger size by the swirling motion of the still-moving core. Some of the smaller fragments ofthe still hot, visco-elastic, disrupting crust acci-dentally got incorporated into the viscous coreand got converted into spherical masses now pre-served as lava balls. Meanwhile, viscous lava atthe £ow front formed a barrier and forced the£ow to thicken with time as hot, less viscouslava continued to be fed towards the front (Sta-siuk et al., 1993). This process ultimately led tothe £attening of the £ow surface towards the me-dial portion and to the steepening of the £owfront at the terminal end.The disposition of crustal fragments across the

Bodshil £ow indicates that slabs were formed notonly at the £ow front but also along the £owmargins. Most slabs, blocks and lava balls weretransported along the surface (Fig. 12D) andrafted over to the terminal front or margins ofthe still-advancing £ow. Long periods of sus-tained shear gradually degraded the blocks (An-derson et al., 1998). Increased shear rates causedattrition of slabs and blocks, grinding their edgesand sides to a pasty mass, which, together withthe passively risen lava, constitute the interslabcrust. Thus, the interslab crust bound the slabstogether and enabled their preservation. Subse-quently, the core halted and stagnated to itspresent form.

8. Discussion

Most simple £ows in the DVP are characterisedby greater thickness (up to 15 m), broad £owfronts and nearly £at upper surfaces capped bydiscontinuous £ow top breccia. The internal struc-ture of most of the simple £ows consists of vesic-ular crust with or without £ow top breccia, thick,non-vesicular, dense (but jointed) core and asharp glassy base. In general, these have been pre-viously considered to be aa £ows. Recent studiesby Managave (2000) and ¢elds checks by us atdi¡erent localities around Pune, Mahad, Mahaba-

Fig. 9. Modal plagioclase vs. glass (A) and mean vesicle de-formation parameter D vs. crystallinity (B) from the Bodshil£ow.

VOLGEO 2539 30-1-03


leshwar, Dhadgaon and Toranmal suggest thatthe simple £ows are quite di¡erent in morphologycompared to aa £ows from Hawaii and other re-gions. The signi¢cant di¡erences include: (a) thelack of prominent £ow bottom breccia, (b) ab-sence of fragments of £ow top breccia in thecore, (c) lack of upper core extending into the£ow top breccia, (d) absence of signi¢cantly de-formed vesicles, and (e) regional extent and sheet-like geometry of the £ows (in contrast to the re-stricted aerial extent and channel structure in aa).If these £ows are considered to be a type of aa, itfollows that they were emplaced at high volumet-ric £ow rates (Rowland and Walker, 1990). How-ever, given the scale and sheer extent of these£ows, it is di⁄cult to envisage how such a ratecould have been maintained throughout the latterhalf of the Deccan eruptions.The sheet-like geometry and the three-tiered

structures of the simple £ows closely resemblethat of typical in£ated pahoehoe, but appear tobe more akin to the distal-type of aa. Moreover,the presence of £ow top breccia and absence of£ow bottom breccia can be accounted for by thefact that the distal-type aa advances without thetypical caterpillar-track motion. It has been sug-gested that these £ows erupted as low-viscositysheet £ows (Managave, 2000) that in£ated consid-erably. However, the subsequent cooling historyof these £ows varied from that of typical pahoe-hoe sheets. The cause and nature of this changehas yet to be determined, although it may be re-lated to increased e¡usion rates and gradients.

Geomorphological studies (Choubey, 1971a)have shown that the surface onto which the Dec-can £ows erupted underwent extrusive peneplana-tion, lateritisation and uplift prior to volcanism inresponse to a starting plume (see Campbell andGri⁄ths, 1990). Uplift took the form of a gentlearch centred on the Pachmari Hills south of theNarmada valley, where the contact between the£ows and the underlying Gondwana sediments isat present approximately 1000 m higher than it isat the margins of the traps (Choubey, 1971b).Rapid northward migration of the Indian plate(at 15 cm/yr) produced an elevated plume tract(volcanic ridge of Hooper, 1990) and may be re-sponsible for the peculiar distribution of com-pound pahoehoe and aa £ows in the DVP (Fig.1). For the estimated Deccan plume head diame-ter of V1000 km the maximum rate of uplift atthe plume axis is predicted to be 20^40 m/Ma anda maximum uplift of 500^1000 m should be at-tained when the plume is at a depth of 100^200km (Campbell and Gri⁄ths, 1990). Thus, onepossible explanation could be that steeper topo-graphic gradients could have existed along theperiphery of the plume tract and could have con-trolled the lava morphology and emplacementmechanism. If this was indeed the case, then thechange in £ow morphology across the stratigra-phy does not necessarily imply an increase in therate of lava supply. Thus, the initial lavas eruptedon gentle slopes and developed into the com-pound pahoehoe of the lower formations. Withgreater plume impact and an increase in the thick-

Table 4Vesicle deformation measurements for samples from the Bodshil £ow

Sr. no. Sample no. Number of vesicles measured Description Vesicle deformation parameter D

1 BOD11 34 pahoehoe crust 0.19502 BOD1/2 38 pahoehoe crust 0.09833 BOD2/2 166 pahoehoe crust 0.16454 BOD3/2 62 pahoehoe crust 0.17365 BOD4/2 239 pahoehoe crust 0.21266 BOD3 101 interslab crust 0.18337 BOD4 42 interslab crust 0.04868 BOD5/2 71 interslab crust 0.12579 BOD12 17 transitional crust 0.180010 BOD14 14 transitional crust 0.200611 BOD13 12 transitional crust 0.5713

VOLGEO 2539 30-1-03


ness of the lava pile, the subsequent lavas belong-ing to the upper formations had to negotiatesteeper gradients and developed into simple types.This can also explain the radial distribution andgreater areal extent of the upper formations, e.g.the Poladpur Formation is exposed from near

Nagpur in the east to Belgaum in the south.Such a model also explains the observed regionaldips in the upper formations and the oversteppingof the formations (Cox, 1983).As mentioned earlier, most simple £ows in the

DVP have characteristics similar to the distal-type

Fig. 10. Sketches of vesicles from hand specimens depicting: A: some deformation in the pahoehoe crust near ‘source’; B: leastdeformation in the interslab crust; and C: most deformed vesicles from the transitional crust.

VOLGEO 2539 30-1-03


aa. If distal-type aa dominates the upper stratig-raphy of the province and if proximal-type aa£ows do not exist on a regional scale in theDVP, the pahoehoe^aa transition in the Bodshil£ow becomes all the more signi¢cant. Such a tran-sition implies that most proximal-type aa £owssensu stricto in the DVP would be of small dimen-sions ^ an observation similar to Hawaii and oth-er provinces. The Bodshil £ow would probablyremain unique amongst £ows in the DVP sincetransitions would be rarely documented becauseof the age, poor surface exposures and greaterdegree of weathering and erosion.

9. Conclusions

Slabby pahoehoe like the spiny, toothpaste andsharkskin varieties are lava types transitional toaa (Rowland and Walker, 1987). Slabby pahoe-hoe forms when the crust completely disrupts dueto a high rate of lava £ow that is too great for thecrust to accommodate the shear strain plastically.In the 1984 Mauna Loa eruption, slabby pahoe-hoe marked the transition to aa within 3^5 kmfrom the vent (Lipman and Banks, 1985, 1987).A similar transition was also observed at a dis-tance of 1.9 km from the vent in the case of a lavachannel at Kalauea Volcano during the May 1997eruption (Cashman et al., 1999). The Bodshil£ow, which is a small £ow (length V1 km),also records such a transition at about 800 mfrom the ‘source’. Thus, most £ows showing mor-phological transformation (pahoehoe^aa transi-tion) deviate from simple correlations betweenlength of lava £ows and eruption rates (Walker,

Table 5Major oxide (wt%) and trace element concentrations (ppm)of samples from the Bodshil £ow

BOD co BOD cr BOD sq Average

SiO2 52.40 56.10 51.80 53.43TiO2 1.25 1.00 1.08 1.11Al2O3 14.70 14.10 15.00 14.60Fe2OT3 12.30 10.70 11.50 11.50MnO 0.24 0.14 0.18 0.19MgO 5.22 4.58 5.12 4.97CaO 9.45 8.32 9.48 9.08Na2O 2.28 2.22 2.25 2.25K2O 0.80 0.89 0.83 0.84P2O5 0.10 0.09 0.09 0.09Total 98.74 98.14 97.33 98.07Trace elementsBa 227.00 245.00 228.00 233.33Be 0.70 0.70 0.60 0.67Cd 0.30 0.30 0.30 0.30Co 44.00 43.00 49.00 45.33Cr 134.00 164.00 148.00 148.67Cs 0.17 0.15 0.14 0.15Cu 120.00 91.00 161.00 124.00Ga 19.00 17.00 19.00 18.33Hf 3.20 2.50 2.80 2.83In 0.06 0.06 0.06 0.06Nb 7.00 5.90 6.30 6.40Ni 62.00 81.00 73.00 72.00Pb 5.00 5.00 5.00 5.00Rb 15.00 23.00 17.00 18.33Sc 36.00 33.00 36.00 35.00Sn 1.20 2.20 0.80 1.40Sr 150.00 159.00 191.00 166.67Ta 0.46 0.40 0.40 0.42Th 3.90 3.20 3.40 3.50Tl 0.08 0.11 0.12 0.10U 0.72 0.51 0.50 0.58V 314.00 174.00 212.00 233.33Zn 101.00 88.00 92.00 93.67Zr 118.00 100.00 107.00 108.33Ce 34.00 29.00 33.00 32.00Dy 4.80 4.40 5.40 4.87Er 2.80 2.50 3.30 2.87Eu 1.30 1.20 1.40 1.30Gd 4.80 4.40 5.60 4.93Ho 1.00 0.90 1.20 1.03La 15.00 15.00 19.00 16.33Lu 0.41 0.36 0.46 0.41Nd 16.00 15.00 19.00 16.67Pr 3.90 3.60 4.50 4.00Sm 4.00 3.60 4.30 3.97Tb 0.82 0.70 0.86 0.79Tm 0.43 0.38 0.49 0.43Y 29.00 29.00 38.00 32.00Yb 2.70 2.30 2.90 2.63

Table 5 (Continued).

BOD co BOD cr BOD sq Average

RatiosTiO2/P2O5 12.50 11.11 12.00 11.89MgO/TiO2 4.18 4.58 4.74 4.48Zr/Nb 16.86 16.95 16.98 16.93Ba/Sr 1.51 1.54 1.19 1.40Ba/Y 7.83 8.45 6.00 7.29Ba/Zr 1.92 2.45 2.13 2.15

Core: BOD co; crust: BOD cr; interslab crust: BOD sq.

VOLGEO 2539 30-1-03


1973). This may be due to the fact that increasedcrustal disruption leads to greater lava immobil-ity. The rate at which a £ow advances and itsthickness are directly related to its thermal struc-ture (Stasiuk et al., 1993). In the case of the Bod-shil £ow, the thickening of the £ow is limited tothe terminal portions, where disrupting crust hasinduced greater cooling. Thickening of the £owfront is attributed to cooling (Stasiuk et al.,1993) and complement ¢eld observations of £owfronts thickening with time in spite of waningeruption rates (Walker, 1967).A key feature of the pahoehoe^aa transition is

that it occurs at higher strain rates for lava with asmaller apparent viscosity or vice versa. Such acondition is consistent with an energy-£ux criteri-on for the failure of solidifying lava. In the case ofthe Bodshil £ow, evidence for increase in bothlava crystallinity and strain rate when the transi-tion occurs con¢rms expectations from earlierindependent studies (e.g. Peterson and Tilling,

1980; Kilburn, 1990; Cashman et al., 1999; Po-lacci et al., 1999).Kilburn (1981) invoked the Bingham-£uid con-

cept to account for the irreversible transition ofpahoehoe to aa (also see Kilburn, 1990). A Bing-ham £uid (Bingham, 1922) is characterised by aviscosity (that is constant) and a yield strength(dy). Unlike a Newtonian £uid where stress is al-ways proportional to the rate of shear, shearstress on Bingham £uid must exceed the yieldstrength to produce a rate of shear greater thanzero. We assume that at least the terminal part ofthe Bodshil £ow was a natural Bingham £uid in alaminar regime. The £ow, therefore, possessed ayield strength, due to which the upper part (crust)disintegrated into slabs and moved along the un-con¢ned £ow.The Bodshil £ow, described in the present pa-

per, is thus unique amongst several £ows from theDVP because it is probably the ¢rst to recordslabby pahoehoe and provides evidence for the

Fig. 11. Concentration of selected major and trace elements normalised to primitive mantle concentrations in the core (BOD co),crust (BOD cr) and interslab crust (BOD sq). The average concentration for all samples from the Bodshil £ow is also plotted.

VOLGEO 2539 30-1-03


incipient transition of basaltic lava from pahoe-hoe to aa.

Acknowledgements

This work was undertaken by R.A.D duringthe tenure of Senior Research Fellowship fromthe Council of Scienti¢c and Industrial Research,New Delhi. G.D and N.B. are thankful to theDepartment of Science and Technology (ES/5(9)/WB/Proj/96) for the ¢nancial support. We thankAnil Sail for assistance during the ¢eldwork. Weare grateful to Margherita Polacci for reprints.Thanks are also due to Dr. Wouter Bleeker, Con-tinental Geoscience Division, Geological Surveyof Canada, for the geochemical analyses. Thispaper bene¢ted greatly from the reviews of Dr.

Christopher Kilburn and Dr. Steven Anderson.We are indebted to them not only for providinginsights into the several aspects of the transitionbut also for the reprints and relevant literature.We would also like to acknowledge Prof. BruceMarsh for his encouragement, patience and kindconsiderations.

References

Anderson, S.W., Stofan, E.R., Plaut, J.J., Crown, D.A., 1998.Block size distributions on silicic lava £ow surfaces: impli-cations for emplacement conditions. Geol. Soc. Am. Bull.110, 1258^1267.

Anderson, S.W., Stofan, E.R., Smrekar, S.E., Guest, J.E.,Wood, B., 1999. Pulsed in£ation of pahoeheo lava £ows:implications for £ood basalt emplacement. Earth Planet.Sci. Lett. 168, 7^18.

Fig. 12. Cartoons depicting stages in the evolution of the Bodshil £ow.

VOLGEO 2539 30-1-03


Aubele, J.C., Crumpler, L.S., Elston, W., 1988. Vesicle zona-tion and vertical structure of basalt £ows. J. Volcanol. Geo-therm. Res. 35, 349^374.

Bean, J.E., Turner, C.A., Hooper, P.R., Subbarao, K.V.,Walsh, J.N., 1986. Stratigraphy, composition and form ofthe Deccan Basalts, Western Ghats, India. Bull. Volcanol.48, 61^83.

Bingham, E.C., 1922. Fluidity and Plasticity. McGraw-Hill,New York, 440 pp.

Bondre, N.R., Dole, G., Phadnis, V.M., Duraiswami, R.A.,Kale, V.S., 2000. In£ated pahoehoe lavas from the San-gamner area of the western Deccan Volcanic Province.Curr. Sci. 78, 1004^1007.

Campbell, I.H., Gri⁄ths, R.W., 1990. Implications of mantleplume structure for the evolution of £ood basalts. EarthPlanet. Sci. Lett. 99, 79^93.

Cashman, K.V., Mangan, M.T., Newmann, S., 1994. Surfacedegassing and modi¢cations to vesicle size distributions inactive basalt £ows. J. Volcanol. Geotherm. Res. 61, 45^68.

Cashman, K.V., Kauahikaua, J.P., Thornber, C., 1997. Cool-ing and crystallization in open lava channels (Abstract).EOS Trans. AGU 78, F793.

Cashman, K.V., Thornber, C., Kauahikaua, J.P., 1999. Cool-ing and crystallisation of lava in open channels, and thetransition of pahoehoe lava to aa. Bull. Volcanol. 61, 306^323.

Choubey, V.D., 1971a. Pre-Deccan Traps topography in cen-tral India and crustal warping in relation to Narmada Riftstructure and volcanic activity. Bull. Volcanol. 35, 660^685.

Choubey, V.D., 1971b. Vertical tectonics of the Satpura basin,Central India. Ann. Geol. Dept. AMU 5^6, 327^352.

Cox, K.G., 1983. The Deccan Traps and the Karoo: strati-graphic implications of possible hot-spot origins. In: IAV-CEI Abstracts, XVII IUGG General Assembly, Hamburg,p. 96.

Cox, K.G., Hawkesworth, C.J., 1985. Geochemical stratigra-phy of the Deccan Traps at Mahabaleshwar, Western Ghats,India, with implications for open system magmatic process-es. J. Petrol. 26, 355^377.

Crisp, J., Cashman, K.V., Bonini, J.A., Hougen, S.B., Pieri,D.C., 1994. Crystallization history of the 1984 Mauna Loalava £ow. J. Geophys. Res. 99, 7177^7198.

Deshmukh, S.S., 1988. Petrographic variations in compound£ows in Deccan Traps and their signi¢cance. In: Subbarao,K.V. (Ed.), Deccan Flood Basalts. Mem. Geol. Soc. India10: 305^319.

Devey, C.W., Lightfoot, P.C., 1986. Volcanological and tec-tonic control of stratigraphy and structure in the westernDeccan Traps. Bull. Volcanol. 48, 195^207.

Dingwell, D.B., Webb, S.L., 1989. Structural relaxation in sil-icate melts and non-Newtonian melt rheology in geologicprocesses. Phys. Chem. Mineral. 16, 508^516.

Duraiswami, R.A., Bondre, N.R., Dole, G., Phadnis, V.M.,Kale, V.S., 2001. Tumuli and associated features from thewestern Deccan Volcanic Province, India. Bull. Volcanol.63, 435^442.

Duraiswami, R.A., Bondre, N.R., Dole, G., Phadnis, V.M.,2002. Morphology and structure of £ow-lobe tumuli fromthe western Deccan Volcanic Province, India. J. Geol. Soc.India 60, 57^65.

Godbole, S.M., Rana, R.S., Natu, S.R., 1996. Lava stratigra-phy of Deccan basalts of western Maharashtra. GondwanaGeol. Mag. Spec. Publ. 2, 125^134.

G.S.I., 1998. Quadrangle Geological Map. Government of In-dia.

Hon, K., Kauahikaua, J., Denlinger, R., MacKay, K., 1994.Emplacement and in£ation of pahoehoe sheet £ows: obser-vations and measurements of active lava £ows on KilaueaVolcano, Hawaii. Geol. Soc. Am. Bull. 106, 351^370.

Hooper, P.R., 1990. The timing of crustal extension and theeruption of continental £ood basalts. Nature 345, 246^249.

Keszthelyi, L., 1994. Calculated e¡ects of vesicles on the ther-mal properties of cooling basaltic lava £ows. J. Volcanol.Geotherm. Res. 63, 257^266.

Keszthelyi, L., 1995. A preliminary thermal budget for lavatubes on the Earth and planets. J. Geophys. Res. 100,20411^20420.

Keszthelyi, L., McEwen, A.S., Thordarson, T., 1999a. Terres-trial analogs and thermal models for Martian £ood lavas(Abstract). Lunar and Planetary Science Conference, vol.30. Houston, TX.

Keszthelyi, L., Self, S., Thodarson, T., 1999b. Application ofrecent studies on the emplacement of basaltic lava £ows tothe Deccan Traps. In: Subbarao, K.V. (Ed.), Deccan Vol-canic Province. Mem. Geol. Soc. India 43: 485^520.

Kilburn, C.R.J., 1981. Pahoehoe and aa lavas: a discussionand continuation of the model by Peterson and Tilling.J. Volcanol. Geotherm. Res. 11, 373^389.

Kilburn, C.R.J., 1990. Surfaces of aa £ow ¢elds on MountEtna, Sicily: morphology, rheology, crystallisation and scal-ing phenomena. In: Fink, J.H. (Ed.), Lava Flows andDomes: Emplacement Mechanisms and Hazard Implica-tions. Springer, Berlin, pp. 129^156.

Kilburn, C.R.J., 1993. Lava crust, aa £ow lengthening and thepahoehoe^aa transition. In: Kilburn, C.R.J., Luongo, G.(Eds.), Active Lava Flows: Monitoring and Modelling.UCL Press, London, pp. 263^280.

Kirkpatrick, R.J., 1976. Towards a kinetic model for the crys-tallization of magma bodies. J. Geophys. Res. 81, 2565^2571.

Kirkpatrick, R.J., 1977. Nucleation and growth of plagioclasein the Hawaiian lava lakes Makaopuhi and Alae. Geol. Soc.Am. Bull. 88, 78^84.

Lipman, P.W., Banks, N.G., 1987. Aa £ow dynamics, MaunaLoa 1984. In: Decker, R.W., Wright, T.L., Stau¡er, P.H.(Eds.), Volcanism in Hawaii, vol. 2. US Geol. Surv. Prof.Pap. 1350: 1527^1567.

Lipman, P.W., Banks, N.G., Rhodes, J.M., 1985. Degassinginduced crystallization of basaltic magma and e¡ects on lavarheology. Nature 317, 604^607.

Macdonald, G.A., 1953. Pahoehoe, aa and block lava. Am.J. Sci. 251, 169^191.

VOLGEO 2539 30-1-03


Managave, S., 2000. The Geology Around Kurundwad. Un-published M.Sc. Dissertation. University of Pune, Pune, In-dia.

Marsh, B.D., 1981. On the crystallinity, probability of occur-rence and rheology of lava and magma. Contrib. Mineral.Petrol. 78, 85^98.

Mitchell, C., Widdowson, M., 1991. A geological map ofsouthern Deccan Traps, India. J. Geol. Soc. London 148,495^505.

Moore, H.J., 1987. Preliminary estimates of the rheologicalproperties of 1984 Mauna Loa lava. In: Decker, R.W.,Wright, T.L., Stau¡er, P.H. (Eds.), Volcanism in Hawaii,vol. 2. US Geol. Surv. Prof. Paper 1350: 1569^1588.

Neilson, R.L., Dungan, M.A., 1983. Low pressure mineral-melt equilibra in natural anhydrous ma¢c systems. Contrib.Mineral. Petrol. 84, 310^326.

Peterson, D.W., Tilling, R.I., 1980. Transition of basaltic lavafrom pahoehoe to aa, Kilauea Volcano, Hawaii: ¢eld obser-vations and key factors. J. Volcanol. Geotherm. Res. 7, 271^293.

Phadke, A.V., Sukhtankar, R.K., 1971. Topographic studies ofDeccan Traps hills around Poona, India. Bull. Volcanol. 35,709^718.

Pinkerton, H., Stevenson, R., 1992. Methods of determiningthe rheological properties of magmas at sub-liquidus tem-peratures. J. Volcanol. Geotherm. Res. 53, 47^66.

Polacci, M., Papale, P., 1997. The evolution of lava £ows fromephemeral vents at Mount Etna: insights from vesicle dis-tribution and morphological studies. J. Volcanol. Geotherm.Res. 76, 1^17.

Polacci, M., Cashman, K.V., Kauahikaua, J.P., 1999. Texturalcharacterization of the pahoehoe^aa transition in Hawaiianbasalt. Bull. Volcanol. 60, 595^609.

Reches, Z., Fink, J., 1988. The mechanism of intrusion of theInyo Dike, Long Valley Caldera, California. J. Geophys.Res. 93, 4321^4334.

Rowland, S.K., Walker, G.P.L., 1987. Toothpaste lava: char-acteristics and origin of a lava-structural type transitionalbetween pahoehoe and aa. Bull. Volcanol. 52, 631^641.

Rowland, S.K., Walker, G.P.L., 1990. Pahoehoe and aa inHawaii: volumetric £ow rate controls the lava structure.Bull. Volcanol. 52, 615^628.

Saar, M., Manga, M., 1999. Permeability porosity relationshipin vesicular basalt. Geophys. Res. Lett. 26, 111^114.

Self, S., Keszthelyi, L., Thordarson, Th., 1998. The importanceof pahoehoe. Annu. Rev. Earth Planet. Sci. 26, 81^110.

Self, S., Thodarson, T., Keszthelyi, L., 1997. Emplacement ofcontinental £ood basalt lava £ows. In: Mahoney, J.J., Cof-¢n, M. (Eds.), Large Igneous Provinces Geophys. Monogr.Ser. AGU, Washington DC, 100: 381^410.

Shaw, H.R., 1969. Rheology of basalt in the melting range.J. Petrol. 10, 510^535.

Stasiuk, M.V., Jaupart, C., Sparks, R.S.J., 1993. In£uence ofcooling on lava £ow dynamics. Geology 21, 335^338.

Stearns, H.T., 1926. The Keaiwa or 1823 lava £ow from Ki-lauea volcano, Hawaii. J. Geol. 34, 336^351.

Stein, D.J., Spera, F.J., 1992. Rheology and microstructure ofmagmatic emulsions: theory and experiments. J. Volcanol.Geotherm. Res. 49, 57^174.

Subbarao, K.V., Bodas, M.S., Hooper, P.R., Walsh, J.N.,1988. Petrogenesis of Jawhar and Igatpuri formations ofwestern Deccan Trap Province. Mem. Geol. Soc. India 10,253^280.

Swanson, D.A., 1973. Pahoehoe £ows from the 1969^1971Mauna Ulu eruption, Kilauea Volcano, Hawaii. Geol. Soc.Am. Bull. 84, 615^626.

Taylor, G.I., 1934. The formation of emulsions in de¢nable¢elds of £ow. Proc. R. Soc. London A 146, 501^523.

Thordarson, T., Self, S., 1998. The Roza Member, ColumbiaRiver Basalt Group: a gigantic pahoehoe lava £ow ¢eldformed by endogeneous processes? J. Geophys. Res. 103(B11), 27411^27445.

Walker, G.P.L., 1967. Thickness and viscosity of Etnean lavas.Nature 213, 484^485.

Walker, G.P.L., 1971. Compound and simple £ows and £oodbasalts. Bull. Volcanol. 35, 579^590.

Walker, G.P.L., 1973. Lengths of lava £ows. Phil. Trans. R.Soc. London A 274, 107^118.

Walker, G.P.L., 1991. Structure, and origin by injection undersurface crust, of tumuli, ‘lava rises’, ‘lava-rise pits’, and ‘lavain£ation clefts’ in Hawaii. Bull. Volcanol. 53, 546^558.

Webb, S.L., Dingwell, D.B., 1990. Non-Newtonian rheologyof igneous melts at high stresses and strain rates: experimen-tal results from rhyolite, andesite, basalt and nephelinite.J. Geophys. Res. 95, 15695^15701.

Wenthworth, C.K., Macdonald, G.A., 1953. Structures andforms of basaltic rocks in Hawaii. US Geol. Surv. Bull.994, 98 pp.

VOLGEO 2539 30-1-03


slabby pahoehoe from the western deccan volcanic province: evidence for incipient pahoehoe^aa...

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