On the conditions of magma mixing and its bearing on andesite production in the crust

ARTICLE

Received 7 Apr 2014 | Accepted 20 Oct 2014 | Published 15 Dec 2014

On the conditions of magma mixing and its bearingon andesite production in the crustMickael Laumonier1, Bruno Scaillet1, Michel Pichavant1, Remi Champallier1, Joan Andujar1 & Laurent Arbaret1

Mixing between magmas is thought to affect a variety of processes, from the growth of

continental crust to the triggering of volcanic eruptions, but its thermophysical viability

remains unclear. Here, by using high-pressure mixing experiments and thermal calculations,

we show that hybridization during single-intrusive events requires injection of high

proportions of the replenishing magma during short periods, producing magmas with

55–58 wt% SiO2 when the mafic end-member is basaltic. High strain rates and gas-rich

conditions may produce more felsic hybrids. The incremental growth of crustal reservoirs

limits the production of hybrids to the waning stage of pluton assembly and to small portions

of it. Large-scale mixing appears to be more efficient at lower crustal conditions, but requires

higher proportions of mafic melt, producing more mafic hybrids than in shallow reservoirs.

Altogether, our results show that hybrid arc magmas correspond to periods of enhanced

magma production at depth.

DOI: 10.1038/ncomms6607

1 ISTO, CNRS-Universite d’Orleans-BRGM, 1a rue de la Ferollerie, Orleans 45071, France. Correspondence and requests for materials should be addressed toB.S. (email: [email protected]).

NATURE COMMUNICATIONS | 5:5607 | DOI: 10.1038/ncomms6607 | www.nature.com/naturecommunications 1

& 2014 Macmillan Publishers Limited. All rights reserved.

mailto:[email protected]

http://www.nature.com/naturecommunications

Although a strong case has been made for andesite arcmagmas being by-products of basalt crystallization1,2, thegeneration of rocks of intermediate compositions has

been also proposed to result from the juxtaposition and large-scale mixing of chemically contrasted magmas3–6, such as basaltand rhyolite3. Whether the latter is a general3,4 or a subsidiary7

process remains, however, debated. In the following, magmamixing is equivalent to hybridization and qualifies ahomogeneous-looking mixture at the hand scale, whereasmingling represents an heterogeneous mixture whose two end-members can be clearly distinguished at the same, or higher,scale.

Typically, magma mixing/mingling occurs during replenish-ment by a mafic magma of a felsic and mushy reservoir8–14, eitherdeep or shallow, which, in the latter case, may trigger a volcaniceruption8. The injection of denser basalt into a lighter reservoirmost likely produces viscous gravity currents spreading at thefloor, leading to a stratified two-layer system15, except whenexcess momentum is available, which may induce fountaining16.Ensuing crystallization can lead to density inversionproducing either local17,18 or wholesale19 overturning andassociated mixing/mingling, depending on viscosity contrast20.Observations and fluid dynamical considerations8–21 have shownthat, in calcalkaline reservoirs, mixing/mingling most likely startseither from such boundary layer instabilities17, with centimetre todecimetre wavelength scale18, or from breakup of injected magmadykes12. These mechanisms readily explain enclave sizes asobserved in plutonic or volcanic rocks18,21, that is, mingledmagmas. However, a further step is required to shift from such amacroscopic heterogeneous rock to a homogeneous mixture atthe hand scale, where mixing is revealed chiefly by the occurrenceof several crystal populations22,23. It is this second step that wewish to address more specifically in our study.

Previous work has considered magma mixing using analogical,numerical or experimental approaches8–20, which have all yieldedvaluable information on the fluid mechanics, leading to efficientstirring. However, the rigorous evaluation of the conditionsfavourable to intimate magma mixing has been hampered by thedifficulty of simulating mixing at high pressure, and by thecomplex fluid-mechanical behaviour of magmas. In particular,the widely held assumption that mixing is feasible at low viscosity,contrasts only9, has not yet received experimental confirmationfor real magmatic mixtures. In addition, early studies9, thoughpaving the way for works such as the present one, did not have inhands precise phase equilibria of magmas under volatile-richconditions, which are vital to constrain thermal exchange duringmixing in arc settings. Laboratory work has concentrated mostlyon mixing between silicate melts of contrasted viscosities11, andthere have been no attempts to explore mixing under high-pressure conditions, which are necessary to dissolve volatiles intosilicate liquids. To date, mixing with crystal-bearing silicate meltshas been explored only at relatively fast strain rates, which aremore appropriate to conduit flow conditions rather than those ofreservoir, and the experiments were done at 1 bar, whichprevented to investigate the role of volatiles24.

We report here viscosity measurements performed duringhigh-pressure felsic–mafic mixing experiments at moderate-to-low strain rates and P-T-fluid conditions relevant to crustalreservoirs25–27. These results are then combined with recentphase equilibria of magmas to evaluate the optimum conditions,leading to efficient mixing, that is, hybrids, when mafic and felsicmagmas come into contact during single-intrusive events. Weconsider both lower and upper crust conditions, as mixingmay occur at virtually all stages of magma evolution, eitherduring emplacement in upper crust28–30, including beforeeruptions31–33, or in the deep crust34–36. In our modelling, we

explore specifically the roles of mafic/felsic mass ratio, shear rateat the felsic–mafic interface and amount of gas, showing thatmixing is promoted by higher values of any of these parameters.We consider our results in the context of the piecemeal nature ofpluton growth37,38, which places stringent limits on the extent ofmixing in shallow reservoirs, and conclude that hybrid andesiteswitness transient increase in mafic magma flux. Altogether, thecalculations are used to infer how variations in magma supplyrate may control magma mixing, hence the geochemical evolutionof magma reservoirs through such a mechanism, and how it mayultimately affect continental crust evolution.

ResultsHigh-pressure mixing experiments. We performed torsionexperiments on felsic–mafic-stacked layers, using an internallyheated deformation apparatus39 at conditions relevant tomagmatic reservoirs in the upper crust (P¼ 300 MPa, T¼ 600–1,200 �C, strain rates ð _gÞ ¼ 3�10� 5 � 3�10� 3s� 1, magmaviscosities¼ 103–1012 Pa s, (refs 25–27)). Before torsionexperiments, we measured the viscosity of felsic and mafic end-members, which were found to agree with model predictions40.The torsion experiments involved fully molten and crystal-bearing felsic end-members and a crystal-bearing mafic magma.By changing the temperature, the crystallinity, hence viscosity, ofthe mafic material changed, so it was either more or less viscousthan the felsic one. We considered that no mixing happenedwhen the stack of layers with straight interfaces was preservedafter deformation (Fig. 1a). In contrast, onset of mixing wascharacterized whenever interface instabilities-yielding mixingfeatures similar to natural textures were observed (Fig. 1b). Thegeometry of the bulk deformation applied in all torsionexperiments is that of simple shear, similar to that, which mayoccur at the vertical interface between a rising blob of magmathrough a more viscous one. In high-temperature runs, however,local departure from simple shear occurred, and a component ofpure shear is probably present in some runs, as might happen ona top part of a rising blob of magma. Owing to technicallimitations, the experiments did not allow us to apply very largeamount of strain (g), in particular at high temperature, whichwould have yielded homogeneous mixtures at the scale of ourexperimental samples (mm to cm) (when conditions appropriateto mixing were reached). Nevertheless, they document theincipient stages of magma mixing at the mm scale,illuminating, in particular, the important transition betweenunmixed and mixed regimes, where single crystals (or small clotsof them) start to be detached from their parent magma andtransported/embedded into foreign melt. We explored thejuxtaposition and deformation of three pairs of magmas (seeSupplementary Fig. 1): (1) dry dacite-haplotonalite, (2) drybasalt-haplotonalite and (3) hydrous basalt-haplotonalite (seeMethods and Supplementary Fig. 1).

The dry dacite-haplotonalite pair was deformed between 850and 1,000 �C (Fig. 2a). In the temperature interval of 850–950 �C,despite the large amount of shear strain applied (g up to 8) and alow viscosity contrast (log(Zhaplotonalite/Zdacite)¼ 0.3), no mixingoccurred for viscosities in the range 1012–108 Pa s. Interfaceinstability appeared only in the experiment conducted at 1,000 �Cwith the boudinage of the layers, corresponding to a measuredviscosity of 107 Pa s. Hence, mixing of viscous and crystal-freemelts, even differing little in viscosity, appears feasible atviscosities lower than 107.5 Pa s under the explored strain rates(Fig. 2a).

The dry basalt-haplotonalite pair is characterized by acontinuous decrease in viscosity contrast as temperatureincreases, owing to the increasing melt fraction of the mafic

ARTICLE NATURE COMMUNICATIONS | DOI: 10.1038/ncomms6607

2 NATURE COMMUNICATIONS | 5:5607 | DOI: 10.1038/ncomms6607 | www.nature.com/naturecommunications



end-member (Fig. 2b). The viscosity contrast reaches itsminimum at a temperature close to 1,170 �C, when both end-members have a viscosity of 104 Pa s (Fig. 2b). Deformationexperiments reveal that this temperature is also the threshold formagma mixing: the initial stack geometry was preserved in the600–1,160 �C temperature interval (Fig. 1a), during which all thestrain was accommodated by the less viscous end-member, thatis, the haplotonalite. We interpret this as the result of the highcrystal fraction in the mafic end-member, which results in acontinuous crystal network that increases the viscosity of thebasalt (Fig. 2b), giving it a yield strength41, as long as the crystalfraction is higher than 46% (volume). At 1,170 �C, that is, only10 �C higher, this touching crystal network collapses, decreasingthe viscosity contrast between both end-members tolog(Zhaplotonalite/Zbasalt)o0.5 (Fig. 2b), yielding mixing featuressimilar to those observed in natural magmas, and which wereobserved in all experiments performed at 1,170 and 1,200 �C(Fig. 1b).

The hydrous basalt-haplotonalite pair shows the effect of wateron magma mixing. Before deformation, both end-members wereannealed under hydrous conditions (see Methods). In thetemperature interval 600–985 �C, corresponding to measuredbulk viscosities between 1012–107.8 Pa s (Fig. 2c), no mixing wasdocumented, all the deformation being accommodated again bythe felsic layer (Fig. 1a, lower panel). Mixing features wereproduced in experiments conducted at 1,000 �C (Fig. 1b, lowerpanel), when the bulk viscosity was around 107.4 Pa s (Fig. 2c),and at higher temperatures. In all cases it occurred when thecrystal fraction was lower than 45%, resulting in a small viscositycontrast (o0.5 log units), as observed under dry conditions(Fig. 2b,c). The role of water appears to be simply to decreasetemperatures, which in the present case is shifted down byB170 �C compared with dry conditions.

The experiments document the fact that the transition betweenthe no-mixing and mixing fields is sharp, occurring over o10 �C

under dry conditions, and over B15 �C with water (Fig. 1), inboth cases corresponding to a transition between locked tounlocked crystal network in the mafic end-member.

Our results therefore reveal that, at shear rates expected tooccur at depth25,26 (o10� 3 s� 1) mixing through enclavedisaggregation is unlikely to occur if the mafic blob has aviscosity higher by 40.3 log unit relative to that of its host.Therefore, the fate of an enclave, that is, whether it disaggregatesvia lateral shearing or remains intact42, will be controlled by theviscosity contrast. The conditions for small viscosity contrast tobe achieved are explored below but our experiments show that theamount of deformation required to disrupt the enclave intosmaller parcels of magma or isolated crystals, is in all cases small(go4); hence, iso-viscous enclaves will disaggregate rapidly ifstrained along with the felsic host. Therefore, mafic enclavespreserved in felsic rocks most probably correspond to the casewhere the mafic blob was immersed into a felsic medium with alower viscosity43. More generally, the experiments evidence thatmagma parcels having crystal contents higher than B50% aresufficiently rigid to resist magmatic shear imposed by thebackground flow field, behaving like solid rather than liquid.The lack of mixing in the first series of deformation experiments(between haplotonalite and dacite melts) at viscosities higher than107 Pa s may be related to the lack of crystals in either magmas.Occurrence of crystals likely affects the local stress field andpossibly enhances mixing by triggering interface instabilities24.

Relative masses of end-members. These rheological constraintsare now used to evaluate how contrasted pairs of natural magmasmay well mix together. We have selected basalt, andesite, daciteand rhyolite compositions whose phase equilibria are known inthe pressure range 0.2–1 GPa44–50. This allows us to establish thetemperature-dependent crystal fraction (Supplementary Fig. 2)and residual liquid composition, and in turn, to calculate the bulk

PP156 - 1,160°C (�loc= 4.3)

Dry

exp

erim

ents

No mixing Mixing

FelsicFelsic Crack

MaficMafic

Jack

et

PP164 - 1,170°C (�loc= 3.6)

FelsicFelsic

Jacket

MaficMafic

MaficMafic

PP296 - 985°C (�loc= 5.1)

MaficMafic

MaficMafic

FelsicFelsic

FelsicFelsic

Jack

et

Hyd

rous

exp

erim

ents

MaficMafic

MaficMaficJacket

FelsicFelsic

FelsicFelsic

PP261 - 1,000°C (�loc= 1.3)

Figure 1 | Illustration of experimental textures. SEM images showing textures obtained when no mixing (a) and mixing (b) happen during torsion

experiments on dry and hydrous samples. The felsic-mafic end-members are separated by a planar interface after torsion at 1,160 �C (or 985 �C for

hydrous conditions), whereas complex textures develop at 1,170 �C (1,000 �C, hydrous) owing to interface instabilities and the collapse of the

touching crystal framework in the mafic layer. Note the small temperature interval (o15 �C) over which such changes occur. Scale bars are 1 mm in PP156,

PP296 and PP261, and 0.5 mm in PP164.

NATURE COMMUNICATIONS | DOI: 10.1038/ncomms6607 ARTICLE




viscosity of each magma at any temperature using an empiricalmodel, which incorporates the effects of abundance and shape ofcrystals (Fig. 3, see Methods).

Using the approach of ref. 9, we calculated the thermal andrheological evolution of a system composed of two juxtaposedmagmas, simulating the replenishment of a felsic reservoir by ahotter more mafic magma, for various fractions of mafic input,exploring a range of initial temperatures for the felsic end-member, whereas initial temperature of the mafic intrusive wasnear its liquidus. Typically, the cooling and associatedcrystallization of hydrous magmas, from basalt to rhyolite,

result in the early increase of its viscosity (Fig. 4). Figure 5shows the cases when the initial felsic temperatures correspond tothat of pre-eruptive storage temperatures, and the mafic end-member is an Mg-rich basalt. In the case of an andesitic reservoir,magma mixing (that is, rocks formed where the felsic and maficcurves cross over or differ by 40.3 log units) happens when themass of injected basalt almost equals the mass of the residentmagma (0.47, Fig. 5a). For more felsic reservoirs, dacitic orrhyolitic, the onset of mixing requires a larger mass fraction,

800750 850 900 950 1,000

Temperature (°C)

DACITE

HTN

Model

Mixed magmasUnmixed magmas

2

4

6

8

10

12Lo

g vi

scos

ity �

900 950 1,000 1,050 1,100 1,150

88

0

0

81

74

64

~100/0

38/0

13/0

46/0

Crystal fractions (mafic/felsic)

Model

Calculatedviscosities�


MAFIC

HTN

Temperature (°C)

PP156PP164

53/38

53/38

45/31

0.38

Cry

stal

fra

ctio

n ef

fect

600 700 800 900 1,000 1,100

Log � ≈ 7 .4 Pa.s

Temperature (°C)


Hydrousfelsic meltHydrous

mafic melt

Maficsuspension

2

4

6

8

10

12

PP296PP261

53/38

100/0

a Dry crystal-free magmas Dry magmas Hydrous magmasb c

Figure 2 | Experimental magma viscosities during mixing. Measured viscosities versus temperature for the three pairs of magmas: dry dacite and

haplotonalite (HTN) (a), dry basalt and haplotonalite (b) and hydrous basalt and haplotonalite (c). Half-filled symbols represent experiments where no

mixing occurs (Fig. 1a), whereas filled symbols are those displaying mixing textures (Fig. 1b). Open symbols represent experiments performed on

single end-members. Dashed red lines in (a) and (b) labelled ‘‘model’’ refer to ref. 40 model applied to end-member compositions. In panels a and b,

solid lines are hand-fitted trends to experimental data. In panel (c), dashed lines are melt viscosities calculated after ref. 40, while blue/red solid lines are

fitted by hand to experimental data. The vertical red arrow shows the calculated effect of 38% crystals. Italicized numbers next to each datum give the

crystal fraction (blue-mafic and red-felsic). Error bars on measured viscosities are shown when larger than symbol size. In panel b, the lowest

viscosities (around 104 Pa s) were beyond reach of instrument capacities, and were calculated (see Methods).

5.0

4.0

3.0

2.0

1.5

1.0

0

1

2

3

4

5

6

7

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7

Crystal fraction (�s)

Log

rela

tive

visc

osity

�re

l (P

a.s)

Figure 3 | Evolution of magma viscosity with crystal fraction and shape

ratio. Each blue curve is labelled with the corresponding shape ratio of

crystals. The calculations have been performed using the viscosity equation

presented in Methods, which allows to incorporate the role of crystal shape,

in addition to their abundance. Experimental magma viscosities obtained

with different crystal shape ratios are shown along with calculated curves:

open square: shape ratio¼ 5; rhombus, shape ratio ¼ 3.5; circle, shape

ratio¼ 4.9; triangle, shape ratio¼ 1.5. Error bars on viscosity

measurements are shown when greater than symbol sizes.

0

1

2

3

4

5

6

7

8

9

10

11

750 800 850 900 950 1,000 1,050 1,100 1,150 1,200

Temperature (°C)

Log

visc

osity

�

Melt viscosity

Bulk viscosity

Upper crust

Lower crust

And Pelée

Bas HAB

Dac Pin. 1

Rhy Usu

Bas Fe

Dac Pin. 2

HPG

Bas Mg St V.

��

Figure 4 | Evolution of magma viscosity with temperature. The melt

fraction trends and compositions needed for viscosity calculations are

derived from the phase equilibrium works36,44–50. Melt viscosities (that is,

the straight lines in the high temperature part of each curve) differ by

o2 orders of magnitude of each other. In contrast, owing to the differences

in their temperature interval of crystallization (the mafic composition

starting their crystallization at higher temperature than the felsic ones),

the viscosity contrast between most felsic–mafic pair (rhyolite-andesite,

andesite-basalt or rhyolite-basalt) is often 43 orders of magnitude,

except at near liquidus conditions. Bas Fe: Fe-rich basalt49; Bas HAB:

high-alumina basalt45; And Pelee: Mt Pelee andesite46; Dac Pin.1: Pinatubo

dacite47; Rhy Usu: Usu Rhyolite50; Bas Mg St V: Mg-rich basalt from

St Vincent36; Dac Pin. 2: Pinatubo dacite48; and HPG: haplogranite44.





Temperature (°C) of the system after replenishment

750

850

900

800

950

1,00

01,

050

1,10

01,

150

1,20

0

0

2

4

6

8

10

Amphibolite (750°C)Basalt (1,210°C)

r = 2.0 r = 3.00.7 GPa 0.56

505560

0 0.2 0.4 0.6 0.8 1

Fraction of injected magma

1,008°C

Hybrid SiO2 wt.%

0

2

4

6

8

10

Greywacke (750°C)Basalt (1,210°C)

r = 2.0

r = 3.0

0.8 GPa 0.58

50556065

0 0.2 0.4 0.6 0.8 1


1,017°C

0

2

4

6

8

10

0 0.2 0.4 0.6 0.8 1


Rhyolite (750°C)Basalt (1,210°C)

1 GPa1,063°C

H2O undersat.

0.58

0.68

505560657075

r = 2.0

r = 3.0

Hybrid SiO2 wt.%Hybrid SiO2 wt.%

Hybrid SiO2 wt.% Hybrid SiO2 wt.%

Hybrid SiO2 wt.%Hybrid SiO2 wt.%

0

2

4

6

8

10

1,067°C1 GPa

Dacite (750°C)Basalt (1,210°C)

H2O undersat.

0 0.2 0.4 0.6 0.8 1


0.69

0.65

50556065

r = 2.0

r = 3.1

0

2

4

6

8

10

Rhyolite (780°C)Dacite (860°C)

r = 3.2

r = 3.5

0.2 GPa 0.42

0 0.2 0.4 0.6 0.8 1


7075

851°C

Log

bulk

vis

cosi

ty

Hybrid SiO2 wt.%


r = 2.9

r = 3.5

0 0.2 0.4 0.6 0.8 1

0.2 GPa

0

2

4

6

8

10


Log

bulk

vis

cosi

ty

0.65

505560657075

1,017°C

Dacite (776°C)Basalt (1,145°C)r = 2.9

r = 3.2

0.2 GPa

0.2 GPa

0

2

4

6

8

10

Log

bulk

vis

cosi

ty

0 0.2 0.4 0.6 0.8 1


0.63

50556065

1,014°C

0

2

4

6

8

10

Log

bulk

vis

cosi

ty

Andesite (890°C)Basalt (1,145°C)

r = 2.9

r = 3.6

0.47

505560

0 0.2 0.4 0.6 0.8 1


1,017°C10

30

50

10

30

10

030

50

10

30

50

20

40

10

0

0

10

30

0

20

50

30

10

40

10

50

10

100

50

30

40

200

10

020

40

50

30

1020

4050

30

10

a e

f

g

h

b

c

d

Figure 5 | Evolution of felsic and mafic magma viscosities with fraction of mafic magma. Viscosities are calculated for an average aspect ratio of crystals,

r (given next to each curve) in each magma (see Fig. 3), and for the case of a hot mafic magma intruding either shallow (a,b,c,d) or deep (e,f,g,h) felsic

reservoirs. Ticks on curves give the wt% of crystals of each end-member. The upper horizontal axis of each panel is graduated in terms of the SiO2 (wt%) of

the mixture corresponding to the mafic fraction. When the viscosity curves cross over, hybrid magmas are produced. Thick vertical grey lines indicate

where the viscosity contrast is o5 (0.3 in log unit). The temperature of each felsic–mafic mixing pair is given by the colour-coded bars. The temperature of

the hybrid corresponding to the minimum amount of mafic magma needed to reach near neutral viscosity contrast is given for each case. The initial

temperature of magmas in each pair is given in parentheses next to composition. The thin-dashed lines extending the basalt curve on panels a,b and c give

the expected bubble effect on viscosity (10% vol). Thick-dashed lines show the viscosity curves of H2O-undersaturated felsic magmas.





respectively of 0.63 and 0.65 (Fig. 5b,c). Similarly, mixing with acolder (crystal-rich and near solidus) felsic end-member increasesthe minimum basalt proportions to beyond 0.7, leading to hotterhybrids (Fig. 6a–c). Using wet basalts as the mafic end-memberincreases the mafic fraction over 0.75 for both the dacite and

rhyolite end-members (Fig. 6d–f). For the common case ofandesite intruding a more felsic reservoir24, similar rheologicalpatterns are observed (Fig. 7), again implying large amounts ofthe more mafic magma for hybridization to proceed. Magmamixing is also sensitive to the mineral assemblage (influencing

Andesite (730°C)Basalt (1,145°C)

r = 2.9

r = 3.2

a

b

c

e

f

d

0

2

4

6

8

10Lo

g bu

lk v

isco

sity

0 0.2 0.4 0.6 0.8 1

Fraction of injected magma0.

71

1,022°C

5560

Hybrid SiO2 wt.%

Hybrid SiO2 wt.%

Hybrid SiO2 wt.% Hybrid SiO2 wt.%

Hybrid SiO2 wt.%

50

30

40

10 0

Dacite (720°C)Basalt (1,145°C)

r = 2.9

r = 3.2

0.2 GPa

0.2 GPa

0

2

4

6

8

10

Log

bulk

vis

cosi

ty

0 0.2 0.4 0.6 0.8 1


0.70

1,018°C

5565 60

020

40

50

30

10


r = 2.9

r = 3.2

0.2 GPa

0

2

4

6

8

10

Log

bulk

vis

cosi

ty

0 0.2 0.4 0.6 0.8 1


0.71

1,022°C

557075 6065

50

30

10

0

20

0

2

4

6

8

10

Andesite (890°C)Basalt HAB (980°C)r = 2.9

r = 3.6

0.2 GPa 0.42

5560

0 0.2 0.4 0.6 0.8 1


Hybrid SiO2 wt.%

928°C

10

20

0

2

4

6

8

10

Dacite (776°C)Basalt HAB (980°C)

r = 3.6

0.2 GPa 0.79

5565 60

0 0.2 0.4 0.6 0.8 1


937°C

r = 2.9

30

020

4050

10

0

2

4

6

8

10

Rhyolite (780°C)Basalt HAB (980°C) r = 2.9

r = 3.2

0.2 GPa 0.76

557075 6065

0 0.2 0.4 0.6 0.8 1


934°C

30

50

10

100


750

850

900

800

950

1,00

01,

050

1,10

01,

150

700

50

30

40

Figure 6 | Evolution of felsic and mafic magma viscosities with fraction of mafic magma. Viscosity evolution of felsic and mafic magmas are

shown for the case of a hot basalt into cold (near solidus) felsic to intermediate reservoirs (a,b,c), and for the case of a cold and wet basalt (HAB, high alumina

basalt) into felsic to intermediate reservoirs (d,e,f). Viscosities are calculated for an average aspect ratio of crystals, r (given next to each curve) in each

magma. Ticks on curves give the wt% of crystals of each end-member. Thick vertical grey lines indicate where the viscosity contrast is o5 (0.3 in log unit).

The temperature of each felsic–mafic mixing pair is given by the colour-coded bars. The temperature of the hybrid corresponding to the minimum amount of

mafic magma needed to reach near neutral viscosity contrast is given for each case. The initial temperature of magmas in each pair is given in parentheses

next to composition. The thin-dashed lines extending the basalt curve on panels a,b and c give the expected bubble effect on viscosity (10% vol).





the shape ratio): for instance, a mafic magma with plagioclase(aspect ratio B4) has a higher viscosity than a similar magmacrystallizing olivine (aspect ratio B2); hence, more mafic melt isneeded in the first instance to reach iso-viscous conditions(B10% more, see Supplementary Fig. 3).

Altogether, apart from when compositionally similar magmasare mixed (that is, injection of a dacite into rhyolite (Fig. 5d), itappears that a significant amount of mafic material (generally450%) is required to reach small viscosity contrast in uppercrustal reservoirs9, the resulting mixtures having SiO2 content inthe range 58–55 wt%. For any given felsic–mafic pair, the viscositycurves of each end-member run parallel to each other beyond thecrossing point, generally differing by less than one log unit. Thissuggests that a range of hybrid compositions can be producedbeyond the crossing point (Figs 5–7), as observed in manyeruptions6,33. The above reasoning has considered only the mafic

into silicic recharge scenario. It has been argued that the reversesituation is common as well5. Progressive arrival of felsic magmasinto a convecting and hotter mafic reservoir corresponds to asituation much more favourable to efficient mixing, as it impliescoexistence of largely (mafic) to fully (silicic) molten end-members (Figs 5–7). Unless there is liquid–liquid immiscibility,such a scenario should end up producing hybrids, providedthat the dynamic conditions of the reservoir are favourable11.In general, any mechanism able to selectively incorporate, forinstance via convective entrainment, small amounts of felsicmelt into a mafic magma (thus lying rightward in all panels ofFigs 5–7) will likely produce hybrids. However, in the commonsituation of a felsic over mafic layers, such a mechanism remainsunclear, as it is the overlying felsic layer, which will convect,thereby entraining underlying mafic parcels. In addition, asfurther detailed below, such a model cannot explain porphyriticandesites having a significant proportion of phenocrysts inheritedfrom the felsic end-member.

As crustal melting is widely believed to occur at the crust–mantle boundary34–36, mixing possibilities were also tested athigh pressures (0.7–1 GPa), exploring two situations: mixingbetween basalt and a discrete felsic layer (Fig. 5e,f), as would arisefrom a compaction-segregation process of a partially meltedcrustal layer35, and mixing between basalt and a partially melted,but unsegregated, metamorphic rock (Fig. 5g,h). For the first case,the mass fraction of mafic material needed is around 0.7, andhybrids so produced have SiO2 contents ranging from 55 wt%(dacite layer, Fig. 5e) to 58 wt% (rhyolite layer, Fig. 5f). In thesecond case, a basalt underplating either amphibolite orgreywacke crust compositions with no melt segregation at750 �C requires a mafic mass fraction of at least 0.58 toproduce a 54–56 wt% SiO2 hybrid (Fig. 5g,h), which issignificantly more viscous, hence less mobile, than when adiscrete felsic layer is involved (104–5 Pa s versus 101–2 Pa s,respectively). Rising the prevailing lower crust temperature to900 �C decreases to 43–45% the amount of basalt required andproduces a 56 wt% SiO2 hybrid (Supplementary Fig. 4).

These calculations thus show that, from the viewpoint of themass of mafic magma needed, mixing is not facilitated by high-pressure conditions, as everything else being equal (compare bwith e and c with f in Fig. 5), mixing at higher pressures requiresslightly higher proportions of mafic end-member than at lowpressure. In other words, hybrid andesites generated in the lowercrust will be, on average, more mafic than those produced in theupper crust. For instance, our calculations show that basalt-rhyolite mixing (Fig. 5c,f) produce basalt-andesite hybrids in thelower crust (SiO2o55 wt%) and true andesites in the upper crust(56 wt%oSiO2o58 wt%). As a potential example, Redoubtvolcano erupted silicic and mafic andesites, and the latter havebeen inferred to have been produced at greater depths than thosemore silicic51. It is also remarkable that most hybrids issued frombasalt-felsic mixing cluster around 55 wt% SiO2, regardless ofwhich felsic end-member (dacite or rhyolite) is being mixed-in.

The calculated T-viscosity conditions for the hybrids (Figs 5–7)fall within the field defined by erupted rocks (Fig. 8), those fromlower crust showing a trend toward higher viscosities than thoseproduced in upper crustal reservoirs. In detail, andesite hybridsproduced from Mg-rich basalt-felsic mixing have temperatures inthe range 1,010–1,060 �C, with magma viscosities around 102–103

Pa s, whereas andesite hybrids arising from mixing betweenwet-differentiated basalt and felsic magmas have somewhat lowertemperatures (920–940 �C), and slightly higher viscosities,102–104 Pa s (Fig. 6b). The crystal content of the former categoryis 20%, all issued from the mafic end-member (Fig. 5). For thecase of wet basalt, the crystal content of the mixture on thermalequilibration is between 20 and 30% (Fig. 6), and apart from


750

850

900

800

950

1,00

01,

050

1,10

01,

150

1,20

0

0

2

4

6

8

10

Log

bulk

vis

cosi

ty

Andesite (1,040°C)

Dacite (780°C)

r = 3.6

r = 3.2

0.2 GPa 0.62

68 66 64

0 0.2 0.4 0.6 0.8 1


Hybrid SiO2 wt.%

Hybrid SiO2 wt.%

940°C

0

2

4

6

8

10

Log

bulk

vis

cosi

ty

Rhyolite (780°C)

Andesite (1,040°C)

r = 3.5

0.2 GPa 0.57

6575 70

0 0.2 0.4 0.6 0.8 1


928°C

r = 3.6

10

30

50

0

20

10

30

50

0

a

b

Figure 7 | Evolution of felsic and intermediate magma viscosities with

fraction of intermediate magma. Viscosities evolutions are shown for the

case of (a) andesite into dacite, and (b) andesite into rhyolite reservoirs.

Viscosities are calculated for an average aspect ratio of crystals, r (given

next to each curve) in each magma. Ticks on curves give the wt% of

crystals of each end-member. Thick vertical grey lines indicate where the

viscosity contrast is o5 (0.3 in log unit). The temperature of each

felsic–mafic mixing pair is given by the colour-coded bars. The temperature

of the hybrid corresponding to the minimum amount of mafic magma

needed to reach near neutral viscosity contrast is given for each case.

The initial temperature of magmas in each pair is given in parentheses

next to composition.





when mixing is between basalt and andesite, they are also onlyof mafic origin. Therefore, an important conclusion, whichemerges from these calculations, is that there should be little,if any, felsic-sourced crystals preserved following mixing(excluding the wet basalt-andesite case). This is particularlytrue for basalt-rhyolite mixing, a common hypothesis for manyporphyritic arc andesites3,4. Hence, a significant part of thecrystal cargo of crystal-rich (40–50%) andesites produced bymixing must form after the andesite condition has been reached,that is, during cooling subsequent to mixing. Surviving felsiccrystals found in hybrids3, in particular quartz and Na-richplagioclase, most probably relate to the sluggish kinetics ofphase dissolution and to the probable causal link between maficintrusions and eruptions8, which arrests mixing phenomena,catching them in the act.

Roles of shear rate and bubbles. So far, we have assumed thatviscosity depends primarily on the crystal load, in addition tocrystal shape and melt composition. Other parameters are knownto influence viscosity, in particular shear rate52 and bubbles53. Toevaluate their role, we use the Pinatubo eruption, a well-knowncase in which magma mixing has been rigorously documented32.Of particular importance in the present context is that, unlike inmost instances, end-members before mixing as well as theresulting mixture, are all well characterized, both in terms ofcomposition and P-T-fluid conditions32,47,54. The andesite hasbeen shown to derive from a mixture of 36% basalt-64% dacite,and looks homogeneous at the hand scale32. Mixing was achievedin the crustal reservoir where the main dacite body was stored.Calculations show that using pre-eruptive conditions of dacite47

and pre-intrusion conditions of the basalt54, 460% basalt isrequired to get similar viscosities between end-members (Fig. 5b),that is, an amount significantly higher than the true mixture.

The simplest way to explain this discrepancy is that theviscosity is affected by either shear rate or bubbles. High shearrates can decrease melt viscosity and promote mixing11,52. Indetail, shear rates associated to enclaves will depend on theirmode of formation. Forceful intrusion of basalt into felsic

reservoirs may give rise to fountaining, which are associated tolocal high shear rates16. Assuming a fountain height of severalhundred metres16, a dyke propagation of 0.1� 1 m s� 1 (ref. 28),the calculated shear rates at the interface between a basalt dykepenetrating a cold dacite reservoir are in the range 10� 3–10� 4 s� 1, that is, comparable to the shear rates needed for theonset of non-Newtonian behaviour (410� 4 s� 1). This may welldecrease the viscosity contrast down to the field where mixing canoperate. In contrast, if enclaves originate from a boundary layerinstability17, then their rise through the felsic layer will bedictated by the mafic–felsic density contrast. Measurements onquenched enclaves have shown that they preserve a densitydifference of about � 600 kg m� 3 with their host17. Using typicalenclaves sizes (cm to dm (refs 18,21)) their ascent rates predictedby the Stokes law are in the range 10� 5–10� 7 m s� 1 (for aviscosity of 105 Pa s, typical of felsic arc magmas27), whichcorrespond to shear rates of 10� 5–10� 8 s� 1, if the rise occursover a 1–10-m thick layer (the approximate thickness of a thermalboundary layer55). Such shear rates will not decrease viscosities,and any viscous blob rising through a static felsic layer will likelykeep its coherency. However, these ascent rates may besuperseded by the vigour of convection in the felsic layer,which is fuelled by the thermal input of the underlying maficintrusion55. Current estimates of convective velocities of felsicreservoirs vary widely12,25, from 10� 8 m s� 1 up to 10� 2 m s� 1,and natural systems quite certainly display a range of convectionrates, depending in particular on magma crystallinity. Clearly,stirring velocities at the high end of this range can provide asignificant impulse to enclave disaggregation42.

Turning to bubbles, recent experiments56 suggest that, undermoderate-to-low shear rates, bubbles lead to a drastic decrease inthe shear viscosity of magmas: a 10% of volume fraction ofbubbles can decrease by up to four orders of magnitude magmaviscosity, and such an effect on mixing is illustrated in Fig. 5. Ifthe basalt replenishing the dacite reservoir coexisted with 10%bubbles (as would occur on its crystallization against the colderdacite17), then the amount of mafic material needed to reachequal viscosities condition is decreased by B10%, that is, closer tothe observed figure.

1

2

3

4

5

6

7

8

9

10

11

12

600 700 800 900 1,000 1,100 1,200

Temperature (°C)

Types of mixing features

Hybrids (32)

Crystals (50)

Layer/banded (6)

Enclaves (13)

b

1

2

3

4

5

6

7

8

9

10

11

12

600 700 800 900 1,000 1,100 1,200

Temperature (°C)

a

Log

visc

osity

�

No mixingMixing in rhyolitesMixing in dacitesMixing in andesitesMixing in basalts

Low crust mix.Up. crust mix

Figure 8 | Viscosities of natural magmas. (a) Comparison between the pre-eruptive viscosity of volcanic products from the literature27,67,68

(squares) and those of hybrids (blue stars) shown on Figs 5–7. Colour code indicates the compositional range. Open squares correspond to erupted

products bearing no evidence of mixing. (b) For those which have such evidence, the type of mixing texture is shown on panel (b) for each erupted

product. The horizontal grey thick line gives the empirical upper viscosity limit beyond which no magma is apparently erupted.





We therefore conclude that both factors (elevated shear ratesand bubbles) are indeed facilitators of mixing in shallowreservoirs, and act to decrease the amount of mafic end-memberneeded to produce hybrids. Accordingly, hybrids resulting frommixing of contrasted magmas (that is, with 410–15 wt%difference in SiO2) and with inferred mafic proportionssignificantly o50%, suggest either elevated water contents (hencebubbles) or high shear rates, both parameters being possiblyinterrelated (that is, high volatile contents decrease density henceincrease ascent velocities). Even so, preservation of felsic crystals(ex quartz) in any significant proportion seems a rather difficultcondition to achieve (see Figs 5 and 6). In addition to kineticseffects, armouring of such minerals by newly grown phases3 mayhelp them withstand dissolution. The role of shear rate andbubbles can be anticipated to be of lesser importance at lowercrustal conditions. In particular, owing to the increase in volatilesolubilities with pressure, bubble effects in lower crustal mixingprocesses are likely to be less important. Only CO2-rich magmasmight be affected, yet the solubility of CO2 in silicate melts at1 GPa is41 wt%, which is higher than dissolved CO2 contentsmeasured in most melt inclusions, in particular in arc settings57.This should further increase the likelihood that hybrids producedin shallow reservoirs are on average more felsic than thoseoriginating in the lower crust.

Role of magma flux. The results above show that to produce anintermediate hybrid rock, a significant amount of mafic magmaneeds to be injected, broadly as much as the mass of the residentmagma (Figs 5–7), which is also similar to that needed to reac-tivate a large reservoir by remobilization (ca 50 vol.% of maficmagma58). It is worth recalling that most of the cases illustratedin Fig. 5 correspond to largely molten felsic reservoirs. Fullyconsolidated or near solidus dacitic/granodioritic or rhyolitic-granitic plutons will require much more mafic input, to achieveconditions appropriate to mixing (480%, Fig. 6). We nowevaluate if such a situation is viable in natural settings.Hybridization, as modelled in Figs 5–7, implicitly assumes thatthe heat exchange occurs between the two end-members only,hence ignores heat loss toward the country rocks. This isequivalent to say that the mafic intrusion is brought rapidly into

contact with the felsic reservoir, or at least over a time periodduring which cooling of the felsic reservoir, owing to conductiveheat transfer toward its surrounding, can be neglected. Analysesat various tectonic settings show that long-term magma fluxes areon the order of 10� 3–10� 4 km3 yr� 1 (ref. 59), althoughconsiderable variations between and within systems aredocumented. Figure 9 shows the time needed for the assemblyof a felsic reservoir in the upper crust, depending on magmafluxes. Also reported are the inferred times of construction forseveral well-known magmatic systems, either plutonic orvolcanic, based on geochronologic and volumetric constraints60.This construction time can be compared with the cooling time ofeach reservoir size (when not fed by mafic input). To estimate thelatter parameter, we use the standard diffusion equation:

t ¼ r2=k

where t is time (s), r is reservoir half size (m) and k thermaldiffusivity (10� 7 m2 s� 1 (ref. 59)). Such an equation assumesthat conduction rather than convection is the main mode of heatdissipation of felsic reservoirs, which on the long-term basis is avalid assumption61. We first assume the limiting condition of aninstantaneous emplacement of a given volume and next considerincremental growth. The characteristic cooling time of a sphericalreservoir is shown superimposed on reservoir construction timeof Fig. 9 (depending on a particular flux rate of magma). For eachflux rate, any reservoir size plotting below the cooling curve is stillfully or partially molten. Hence, if intercepted by an ascendingmafic batch, it will be prone to hybrid production, provided thesupplied amount of mafic magma is broadly equivalent to that ofthe resident magma (or the part, which is still molten), inaccordance with constraints from Figs 5–7. Conversely, anysystem lying above the cooling curve will be essentially solid, andunable to mix with mafic forerunners. For an average magma fluxrate of 10� 3 km3 yr� 1, system volumes lower than 1,000 km3

remain liquid, whereas for a flux rate of 10� 4 km3 yr� 1, thecritical volume beyond which the reservoir will be frozen issignificantly smaller, o1 km3. The geological trend crosses thecooling curve at the high end, which would suggest that mixingoperates in almost all circumstances, with the implication thatlarge volumes of hybrids can be achieved by mixing in upper

10–4 km

3 yr–1

10–3 km

3 yr–1

10–1 km

3 yr–1

Accretion rate: 100 km

3 yr–1

10–7m

2 s–1

–

–4

Dinkey creek Soccoro

Mt Givens

Geyser plut.

Manaslu

Aleutian islands

Santorini

PeléePelée

100

101

102

103

104

105

106

107

108

10–1

Tim

e (y

)

1 10 100 1,000 10,000 100,0000.1

Volume (km3)

Plutons, convergent settingsPlutons, other settingsVolcanoes

Mix. No mix.

10–6m

2 s–1

Figure 9 | Relationships between volumes of reservoir and cooling times. The construction times of intrusive bodies in upper crust60 are compared with

reservoir solidification times with volume (oblique thick dark curve calculated following59, whereas the lighter-shaded field gives cooling times for two

different thermal diffusivities calculated as explained in the main text). Dashed lines give the relationship between average magma flux and reservoir

volume. Horizontal shaded field gives the cooling times for dyke thicknesses of 50 and 200 m. The active (molten) part of Mt Pelee reservoir is shown also

to the left of full reservoir size69, connected to it by a dashed line.





crust. However, for the more likely scenario of incrementalgrowth of plutons37,38, the cooling pattern is dictated by the sizeof sill intrusion as well as by the time interval between twosuccessive magma additions. In such a case, numericalsimulations37,38 have shown that at flux rates lower than10� 3–10� 2 km3 yr� 1, no molten reservoir of any significantsize can develop, which considerably restricts the likelihood ofwidespread efficient mixing in upper crust. Each individual sill(thicknesses of 50–200 m) will cool down solidus in o500–5,000years, which truncates the natural trend at reservoir volumes ofabout 10–100 km3 (Fig. 9). At an average rate of 10� 4 km3 yr� 1,incubation time needed to start accumulating melt in mid-sizedreservoirs (10–100 km3) are in the range 60–120 k yr (ref. 37),beyond which only 10% of the intruded mass remains active (thatis, partly molten, see the case for Mt Pelee Fig. 9). This indicatesthat, in upper crust, mixing will potentially have a petrogeneticrole during the late stage of pluton construction, and affect asmall part only of the whole intrusion. Hence, large reservoirsbuilt in upper crust at average flux rates lower than10� 2–10� 3 km3 yr� 1 will be unfavourable locii to theproduction of hybrid rocks in any significant amounts, insteadpreserving abundant evidence of magma blending andincomplete mixing28–30 (that is, mingled magmas). This rule,however, may be violated whenever a surge in magma-feedingrates happens during batholith construction37,62. Note thatmixing needs not to involve all the mass of felsic magma.Beside the above thermal restrictions, the fluid dynamicsassociated to reservoir recharge may restrict efficient mixing toa small portion of the intruded reservoir as illustrated by thePinatubo eruption63.

DiscussionThe previous arguments show that hybrids require above all vastamounts of mafic magmas (relative to felsic) being injectedrapidly into partially molten felsic reservoirs. Therefore, whenrooted into magmatic systems of inferred small dimensions,emission of abundant hybrid andesites bears witness of enhancedmagmatic production along with that particular mantle–crustsection. Systems characterized by relatively low magma fluxeswill, in contrast, produce bimodal magmatic provinces, if mixingis the main driver of magma compositional diversity. Recentmodelling studies have concluded that magmatic bimodality isfavoured by volatiles-poor magmas and low-heat input36. Thepresent work reinforces and adds a rheological dimension to sucha paradigm, showing that low-heat input and bubble-free regimeinhibit efficient small-scale homogenization, thus preservingmagma diversity.

From a volcanological perspective, the demonstration thatmost crystals of the felsic end-member should melt out on mixingprovides a robust time scale on mixing phenomema in eruptioncontexts. Preservation of felsic phenocrysts shows that thermalexcursions associated to magma intrusion are highly transient.For instance, experiments on Pinatubo dacite show that completemelting of H2O-rich dacite at 930 �C takes place in a day or so47,which provides an upper bound on the time spent between basaltintrusion and andesite eruption for this particular case.

The fact that mixing is not enhanced at high pressures suggeststhat magmas produced in the deep crust may well preserve itschemical heterogeneity. Intermediate to felsic magmas derivedfrom lower crustal processes are thus likely to be faithful mirrorsof lower crustal lithologies at regional scales at least7. However,any sizeable mass of rock, demonstrably hybrid in origin, musthave been produced in the deep crust before emplacement inupper crust, owing to the severe thermal restrictions imposed onmixing efficiency by the mechanism of incremental growth at

shallow levels. At last, the pressure effect has also implications forthe understanding of the chemical evolution of the continentalcrust. A thin crust frequently intruded by hot basalt, as couldhave been the case in the early Earth, has the potential to producelarge amounts of relatively felsic hybrids. Conversely, an increasein crust thickness, correlated to a lower mantle heat flux, arebound to produce less, but more mafic, hybrids, and thus drivethe bulk composition of the continental crust toward a moremafic composition. This may have helped to produce an Archeancontinental crust more silicic than the Phanerozoic one64, as thereare no reasons to believe that mixing was a less importantpetrogenetic mechanism in the Archean than nowadays.

MethodsStarting materials. Samples for torsion experiments were composed of pairsof magmas from the three starting materials whose compositions are given inSupplementary Table 1. The most felsic end-member is a synthetic haplotonalitemanufactured by Schott A.G. (Germany), which has been extensively used inprevious deformation experiments65. Its composition allows the crystallization ofplagioclase BAn30 only over a large range of crystal fraction (0–65% in volume)and temperature (B700–1,220 �C) depending on water content. In the first seriesof experiments, the magma juxtaposed to the haplotonalite was a natural dacite ofSantorini Volcano (Greece), which was previously melted at 1 bar at 1,500 �C toproduce a crystal-free dacite glass. The mafic end-member of the second and thirdseries of experiments was a basalt from a lava flow from Santorini Volcano: thisstarting basalt is holocrystalline and was used directly in the torsion experiments.

Syntheses of starting material. In the first series of experiments, both compo-sitions (haplotonalite and dacite) were melted at 1,500 �C for 3 h in a 1-atm furnaceto get dry, crystal-free, magmas. The water content in both melts is lower than0.1 wt% and no crystals were detected. After the torsion experiments, both magmasremained free of crystals, except the dacite in the experiment conducted at 1,000 �C(see below), which displayed a 20-mm thick crystallized halo at the interface withthe haplotonalite. The second series of experiments involved the crystal-bearingbasaltic magma, together with the same dry, crystal-free, haplotonalite used in thefirst series. Both magmas were used dry in an effort to limit quench phases. Thehaplotonalitic magma remained crystal-free at To1,150 �C, but started to crys-tallize at the highest temperatures investigated (1,150–1,200 �C), producing a thinhalo at the interface with basalt or capsule (o150mm) for the longest experimentaldurations (B8 h). The crystallization interval of the basalt ranges between B1,000and 1,230 �C and the run temperature was adjusted to produce different viscositycontrasts between end-members. See Laumonier et al.66 for a detailed descriptionof the textures and chemical compositions of the phases in the starting materialsand experimental products. The third series of experiments used hydrous crystal-bearing haplotonalite and basalt glasses, which were obtained by a firstcrystallization experiment at P¼ 300 MPa, T¼ 950/1,000 �C, t¼ 7 days in aninternally heated pressure vessel. This first stage produced water-saturated crystal-bearing melts with 38 and 53% (950 �C) and 31 and 45% (1,000 �C) of crystals forthe felsic and the mafic components, respectively. The syntheses performed at950 �C were used in torsion experiments conducted at 600, 715, 950, 975 �C,whereas that at 1,000 �C was used in torsion experiments performed at and 985,1,000, 1,020 �C. The crystal contents of first-step syntheses were preserved in thetorsion experiments conducted at 600 and 715 �C. For other deformationexperiments, small changes in crystal content occurred, toward either higher(that is, a torsion experiment at 985 �C using a synthesis at 1,000 �C) or lowerproportions depending on whether the temperature of mixing experiment washigher or lower than that of the first-step synthesis temperature. These smallchanges in phase proportions were taken into account in the viscosity calculations,using a melt fraction versus temperature relationships calibrated experimentally foreach composition.

Deformation experiments. Torsion experiments were conducted in a Patersonapparatus39 at 300 MPa, strain rates ranging from 3� 10� 5 to 3� 10� 3 s-1, andvarious temperatures (Supplementary Tables 2–5). During torsion experiments, thetorque of the assembly containing the sample was recorded. Then, the torque wascorrected from the jacket contribution to calculate the shear stress and the viscosityof the sample only39. The error on viscosity calculation depends on the intrinsicprecision of the load cell recording the internal torque, on the regularity of theplateau that defines the value of the torque and on the jacket contribution. Themaximal theoretical error on the viscosity was calculated by adding all three factorsand is listed in Tables 2–5 and in Fig. 2 (see main text). Two kinds of samples wereprepared depending on the target of the experiment. To characterize therheological evolution of end-members with temperature, the cylindrical sampleinserted into the Paterson press was composed of a single layer of that end-member(dry, crystal-free dacite and haplotonalite and dry, melt-bearing basalt) whereas toinvestigate the mixing capacities between two magmas, the samples consisted in a





stack of 2–4 layers alternating in composition (Supplementary Fig. 1; see alsoref. 66). When a torsion deformation is applied to either a two- or four-layers stack(Supplementary Fig. 1), it imparts simple shear along the various interfacesbetween compositionally different layers (if their rheological properties differ). Thisshearing mechanism is similar, for instance, to that occurring at the interfacebetween a rising basaltic dyke into a felsic reservoir, or between a basaltic sillflowing at the base of a felsic reservoir and its felsic host.

Viscosity calculations. In addition to direct viscosity measurements performedwith the Paterson press, we calculated the expected magma viscosities at our P-Tconditions. Melt viscosity Zmelt was calculated using the model of ref. 40. The effectof crystals, that is, the relative viscosity Zrel, depends on both the crystal fraction(fS) and the shape ratio (r) of the crystals. In this study, we derived an empiricalmodel to calculate the relative viscosity Zrel, by combining several models andpublished (see Supplementary Table 6) and unpublished data (Fig. 3). This modelis restricted to a crystal fraction lower than 0.6, and assumes that several crystalpopulations with different shape ratios can be considered as a unique populationwith an average shape ratio r. The empirical equation fitting the available data is:

LogZrel ¼a

1þ b expj�fsþ d

1þ e expz�fs

with a¼ 0.373� rþ 4.89, b¼ 17.241� exp0.748 � r, j¼ � 3.96� r, d¼ 0.108�r� 0.292, e¼ 2� 10� 4 exp3.50 � r, z¼ � 24� rþ 52.6.

Injected fraction of mafic magma and viscosity. Input parameters used tocalculate the final temperature reached by a system composed of a felsic reservoirreplenished by a mafic intrusion are given in Supplementary Tables 7,8. Thecalculations were done using the method of ref. 9. The initial temperature of eachmafic magmas was its liquidus temperature. The initial temperatures of the felsicend-members were chosen to encompass the pre-eruptive temperatures of felsicreservoirs in arc settings27, except for the case of the melting of lower crustlithologies (see Supplementary Table 8).

References1. Gill, J. B. Orogenic Andesites and Plate Tectonics Vol. 390 (Springer-Verlag,

1981).2. Pichavant, M., Martel, C., Bourdier, J. L. & Scaillet, B. Physical conditions,

structure, and dynamics of a zoned magma chamber: Mount Pelee (Martinique,Lesser Antilles Arc). J. Geophys. Res. 107, ECV-1 (2002).

3. Eichelberger, J. C. Andesitic volcanism and crustal evolution. Nature 275,21–27 (1978).

4. Reubi, O. & Blundy, J. A dearth of intermediate melts at subduction zonevolcanoes and the petrogenesis of arc andesites. Nature 461, 1269–1273 (2009).

5. Eichelberger, J. C., Chertkoff, D. G., Dreher, S. T. & Nye, C. J. Magmas incollision: rethinking chemical zonation in silicic magmas. Geology 28, 603–606(2000).

6. Ruprecht, P., Bergantz, G. W., Cooper, K. M. & Hildreth, W. The crustalmagma storage system of Volcan Quizapu, Chile, and the effects of magmamixing on magma diversity. J. Petrol. 53, 801–840 (2012).

7. Clemens, J. D. & Stevens, G. What controls chemical variation in graniticmagmas? Lithos 134, 317–329 (2012).

8. Sparks, R. S. J., Sigurdsson, H. & Wilson, L. Magma mixing: a mechanism fortriggering acid explosive eruptions. Nature 267, 315–318 (1977).

9. Sparks, R. S. J. & Marshall, L. A. Thermal and mechanical constraints onmixing between mafic and silicic magmas. J. Volcanol. Geoth. Res. 29, 99–124(1986).

10. Jellinek, A. M. & Kerr, R. C. Mixing and compositional stratification producedby natural convection: 2. Applications to the differentiation of basaltic andsilicic magma chambers and komatiite lava flows. J. Geophys. Res. 104,7203–7218 (1999).

11. De Campos, C. P., Perugini, D., Ertel-Ingrisch, W., Dingwell, D. B. & Poli, G.Enhancement of magma mixing efficiency by chaotic dynamics: anexperimental study. Contrib. Mineral. Petrol. 161, 863–881 (2011).

12. Hodge, K. F., Carazzo, G. & Jellinek, A. M. Experimental constraints on thedeformation and breakup of injected magma. Earth Planet. Sci. Lett. 325, 52–62(2012).

13. Burgisser, A. & Bergantz, G. W. A rapid mechanism to remobilize andhomogenize highly crystalline magma bodies. Nature 471, 212–215 (2011).

14. Andrews, B. J. & Manga, M. Thermal and rheological controls on theformations of mafic enclaves and banded pumices. Contrib. Mineral. Petrol.167, 961 (2014).

15. Snyder, D. & Tait, S. Replenishment of magma chambers: comparison offluid-mechanic experiments with field relations. Contrib. Mineral. Petrol. 122,230–240 (1995).

16. Campbell, I. H. & Turner, J. S. The influence of viscosity on fountains inmagma chambers. J. Petrol. 27, 1–30 (1986).

17. Eichelberger, J. C. Vesiculation of mafic magma during replenishment of silicicmagma reservoirs. Nature 288, 446–450 (1980).

18. Thomas, N. & Tait, S. R. The dimensions of magmatic inclusions as a constrainton the physical mechanism of mixing. J. Volcanol. Geoth. Res 75, 167–178(1997).

19. Huppert, H. E., Sparks, R. S. J. & Turner, J. S. Effects of volatiles on mixing incalc-alkaline magma systems. Nature 297, 554–557 (1982).

20. Huppert, H. E., Stephen, R., Sparks, J. & Turner, J. S. Some effects of viscosityon the dynamics of replenished magma chambers. J. Geophys. Res. 89,6857–6877 (1984).

21. Bacon, C. R. Magmatic inclusions in silicic and intermediate volcanic rocks.J. Geophys. Res. 91, 6091–6112 (1986).

22. Clynne, M. A. A complex magma mixing origin for rocks erupted in 1915,Lassen Peak, California. J. Petrol. 40, 105–132 (1999).

23. Martel, C., Ali, A. R., Poussineau, S., Gourgaud, A. & Pichavant, M.Basalt-inherited microlites in silicic magmas: evidence from Mount Pelee(Martinique, French West Indies). Geology 34, 905–908 (2006).

24. Kouchi, A. & Sunagawa, I. A model for mixing basaltic and dacitic magmasas deduced from experimental data. Contrib. Mineral. Petrol. 89, 17–23(1985).

25. Spera, F. J., Borgia, A., Strimple, J. & Feigenson, M. Rheology of melts andmagmatic suspensions: 1. Design and calibration of concentric cylinderviscometer with application to rhyolitic magma. J. Geophys. Res. 93,10273–10294 (1988).

26. Albertz, M., Paterson, S. R. & Okaya, D. Fast strain rates during plutonemplacement: Magmatically folded leucocratic dikes in aureoles of the MountStuart Batholith, Washington, and the Tuolumne Intrusive Suite, California.Geol. Soc. Am. Bull. 117, 450–465 (2005).

27. Scaillet, B., Holtz, F. & Pichavant, M. Phase equilibrium constraints on theviscosity of silicic magmas: 1. Volcanic-plutonic comparison. J. Geophys. Res.103, 27257–27266 (1998).

28. Paterson, S. R., Pignotta, G. S. & Vernon, R. H. The significance ofmicrogranitoid enclave shapes and orientations. J. Struct. Geol 26, 1465–1481(2004).

29. Perugini, D. & Poli, G. The mixing of magmas in plutonic and volcanicenvironments: analogies and differences. Lithos 153, 261–277 (2012).

30. Caricchi, L., Annen, C., Rust, A. & Blundy, J. Insights into the mechanisms andtimescales of pluton assembly from deformation patterns of mafic enclaves.J. Geophys. Res. 117, B11206 (2012).

31. Cioni, R. et al. Compositional layering and syn-eruptive mixing of aperiodically refilled shallow magma chamber: the AD 79 Plinian eruption ofVesuvius. J. Petrol. 36, 739–776 (1995).

32. Pallister, J. S. et al. Magma mixing at Mount Pinatubo volcano: petrographicaland chemical evidence from the 1991 deposits In: Newhall, C.G., &Punongbayan, R.S. (eds.). Fire and Mud. Eruptions and Lahars of MountPinatubo, Philippines, University of Washington Press, pp 687–732.

33. Kent, A. J., Darr, C., Koleszar, A. M., Salisbury, M. J. & Cooper, K. M.Preferential eruption of andesitic magmas through recharge filtering. NatureGeosci 3, 631–636 (2010).

34. Hildreth, W. & Moorbath, S. Crustal contributions to arc magmatism in theAndes of central Chile. Contrib. Mineral. Petrol. 98, 455–489 (1988).

35. Solano, J. M. S., Jackson, M. D., Sparks, R. S. J., Blundy, J. D. & Annen, C. Meltsegregation in deep crustal hot zones: a mechanism for chemical differentiation,crustal assimilation and the formation of evolved magmas. J. Petrol. 53,1999–2026 (2012).

36. Melekhova, E., Annen, C. & Blundy, J. Compositional gaps in igneous rocksuites controlled by magma system heat and water content. Nature Geosci. 6,385–390 (2013).

37. Schopa, A. & Annen, C. The effects of magma flux variations on the formationand lifetime of large silicic magma chambers. J. Geophys. Res. 118, 926–942(2013).

38. Gelman, S. E., Gutierrez, F. J. & Bachmann, O. On the longevity of large uppercrustal silicic magma reservoirs. Geology 41, 759–762 (2013).

39. Paterson, M. S. & Olgaard, D. L. Rock deformation tests to large shear strains intorsion. J. Struct. Geol. 22, 1341–1358 (2000).

40. Giordano, D., Russell, J. K. & Dingwell, D. B. Viscosity of magmatic liquids: amodel. Earth Planet Sci. Lett. 271, 123–134 (2008).

41. Philpotts, A. R., Shi, J. & Brustman, C. Role of plagioclase crystal chains in thedifferentiation of partly crystallized basaltic magma. Nature 395, 343–346(1998).

42. Blake, S. & Fink, J. H. On the deformation and freezing of enclaves duringmagma mixing. J. Volcanol. Geoth. Res. 95, 1–8 (2000).

43. Scaillet, B., Whittington, A., Martel, C., Pichavant, M. & Holtz, F. Phaseequilibrium constraints on the viscosity of silicic magmas II: implications formafic—silicic mixing processes. Trans. R. Soc. Edinb. 91, 61–72 (2000).

44. Johannes, W. & Holtz, F. Petrogenesis and Experimental Petrology of GraniticRocks Vol. 335 (Springer, 1996).

45. Sisson, T. W. & Grove, T. L. Experimental investigations of the role of H2O incalc-alkaline differentiation and subduction zone magmatism. Contrib. Mineral.Petrol. 113, 143–166 (1993).





46. Martel, C. et al. Effects of fO2 and H2O on andesite phase relations between 2and 4 kbar. J. Geophys. Res. 104, 29453–29470 (1999).

47. Scaillet, B. & Evans, B. W. The 15 June 1991 eruption of Mount Pinatubo. I.Phase equilibria and pre-eruption P–T–fO2–fH2O conditions of the dacitemagma. J. Petrol. 40, 381–411 (1999).

48. Prouteau, G. & Scaillet, B. Experimental constraints on the origin of the 1991Pinatubo dacite. J. Petrol. 44, 2203–2241 (2003).

49. Botcharnikov, R. E., Almeev, R. R., Koepke, J. & Holtz, F. Phase relationsand liquid lines of descent in hydrous ferrobasalt—implications for theSkaergaard intrusion and Columbia River flood basalts. J. Petrol. 49, 1687–1727(2008).

50. Tomiya, A., Takahashi, E., Furukawa, N. & Suzuki, T. Depth and evolution of asilicic magma chamber: Melting experiments on a low-K rhyolite from Usuvolcano, Japan. J. Petrol. 51, 1333–1354 (2010).

51. Coombs, M. L. et al. Andesites of the 2009 eruption of Redoubt Volcano,Alaska. J. Volcanol. Geoth. Res. 259, 349–372 (2013).

52. Dingwell, D. B. & Webb, S. L. Relaxation in silicate melts. Eur. J. Mineral 4,427–449 (1990).

53. Mader, H. M., Llewellin, E. W. & Mueller, S. P. The rheology of two-phasemagmas: A review and analysis. J. Volcanol. Geoth. Res. 257, 135–158 (2013).

54. De Hoog, J. C. M., Hattori, K. H. & Hoblitt, R. P. Oxidized sulfur-rich maficmagma at Mount Pinatubo, Philippines. Contrib. Mineral. Petrol. 146, 750–761(2004).

55. Snyder, D. Thermal effects of the intrusion of basaltic magma into a more silicicmagma chamber and implications for eruption triggering. Earth Planet. Sci.Lett. 175, 257–273 (2000).

56. Pistone, M., Caricchi, L., Ulmer, P., Reusser, E. & Ardia, P. Rheology of volatile-bearing crystal mushes: Mobilization vs. viscous death. Chem. Geol. 345, 16–39(2013).

57. Blundy, J., Cashman, K. V., Rust, A. & Witham, F. A case for CO2-rich arcmagmas. Earth Planet. Sci. Lett. 290, 289–301 (2010).

58. Huber, C., Bachmann, O. & Dufek, J. The limitations of melting on thereactivation of silicic mushes. J. Volcanol. Geoth. Res. 195, 97–105 (2010).

59. White, S. M., Crisp, J. A. & Spera, F. J. Long-term volumetric eruption rates andmagma budgets. Geochem. Geophys. Geosyst. 7, Q03010 (2006).

60. de Saint Blanquat, M. et al. Multiscale magmatic cyclicity, duration of plutonconstruction, and the paradoxical relationship between tectonism andplutonism in continental arcs. Tectonophys 500, 20–33 (2011).

61. Koyaguchi, T. & Kaneko, K. Thermal evolution of silicic magma chambers afterbasalt replenishments. Trans. R. Soc. Edinb. 91, 47–60 (2000).

62. Druitt, T. H., Costa, F., Deloule, E., Dungan, M. & Scaillet, B. Decadal tomonthly timescales of magma transfer and reservoir growth at a calderavolcano. Nature 482, 77–80 (2012).

63. Snyder, D. & Tait, S. Magma mixing by convective entrainment. Nature 379,529–531 (1996).

64. Rollinson, H. Secular evolution of the continental crust: Implications for crustevolution models. Geochem. Geophys. Geosyst. 9, Q12010 (2008).

65. Picard, D., Arbaret, L., Pichavant, M., Champallier, R. & Launeau, P. Therheological transition in plagioclase-bearing magmas. J. Geophys. Res. 118,1363–1377 (2013).

66. Laumonier, M., Scaillet, B. & Arbaret, L. Champallier. Experimental simulationof magma mixing at high pressure. Lithos 196, 281–300 (2014).

67. Takeuchi, S. Preeruptive magma viscosity: an important measure of magmaeruptibility. J. Geophys. Res 116, B10201 (2011).

68. Andujar, J. & Scaillet, B. Relationships between pre-eruptive conditionsand eruptive styles of phonolite–trachyte magmas. Lithos 152, 122–131(2012).

69. Annen, C., Pichavant, M., Bachmann, O. & Burgisser, A. Conditions for thegrowth of a long-lived shallow crustal magma chamber below Mount Peleevolcano (Martinique, Lesser Antilles Arc). J. Geophys. Res. 113, B07209 (2008).

AcknowledgementsThis work is part of the PhD study of ML, and has been supported by theANR-STOMIXSAN, Equipex PLANEX and ERC #279790 projects. Discussions withLuca Caricchi, Fidel Costa, and Tim Druitt on magma mixing processes are gratefullyacknowledged.

Author contributionsM.L. performed the experiments with the help of R.C., and did the calculations. M.L. andB.S. wrote the first draft to which all remaining authors contributed equally.

Additional informationSupplementary Information accompanies this paper at http://www.nature.com/naturecommunications

Competing financial interests: The authors declare no competing financial interests.

Reprints and permission information is available online at http://npg.nature.com/reprintsandpermissions/

How to cite this article: Laumonier, M. et al. On the conditions of magma mixingand its bearing on andesite production in the crust. Nat. Commun. 5:5607doi: 10.1038/ncomms6607 (2014).






http://www.nature.com/reprints/index.html

http://www.nature.com/reprints/index.html


1

Supplementary Figure 1 | Experimental sample set up. Sketches of samples used for mixing experiments

consisting in the vertical juxtaposition of 2 and 4 layers simulating simple shear deformation by twisting, and

allowing strain partitioning between end-members. SEM images presented in Fig. 1 (main text) are tangential

sections of different experiments.

2

Supplementary Figure 2 | Crystal fraction and magma viscosity. (a) Evolution of the crystal fraction Φs with

temperature for each magma composition used in the calculations, expressed as the difference relative to the

liquidus temperature. (b) Average shape ratio of the crystal suspension in the different magmas used for viscosity

calculations. The grey shaded area below Φs = 0.2 highlights the domain where crystal fraction has little effect on

the relative viscosity. At Φs > 0.2, most of the magmas have an average crystal shape ratio between 3.0 and 3.6,

except the feldspar-poor basalt from Saint Vincent (Bas Mg). Numbers next to each curve indicate the average

shape ratio used to calculate the viscosity of each magma.

3

Supplementary Figure 3 | Effect of the shape ratio of crystals on mixing. The shape ratio influences the

relative viscosity and consequently the bulk viscosity, hence mixing. While the average shape ratio varies little

over the crystal fraction 0.2 – 0.6, small differences in composition of otherwise similar magmas may affect their

average crystal shape ratio, and consequently their respective bulk viscosity (compare Bas Fe and Bas Mg on Fig.

4 in main text). The basalt from Skaergaard intrusion1 is mostly composed of clinopyroxene, plagioclase and

magnetite, and has an intermediate average shape ratio. As shown, a lower shape ratio (such as with an

assemblage dominated by olivine) decreases the amount of injected mafic magma fraction required to reach

favourable conditions for magma mixing, with a mixing ratio of 0.47. In contrast, if the crystallising phase

assemblage is dominated by plagioclase (larger shape ratio) a larger fraction of a mafic magma (0.58) is necessary

to produce a hybrid magma.

4

Supplementary Figure 4 | Effect of host rock temperature on mixing in the lower crust. The results are

shown for hybridization of (a) basalt - greywacke and, (b) basalt - amphibolite. The hatched vertical fields

correspond to mixing conditions when the crust is at 750°C (shown on Fig 5). Note that for the amphibolite case,

equal viscosities are reached at crystal contents beyond 50%, ie outside the calibrated range of the viscosity

equation used here, and mixing is unlikely in this case.

5

Supplementary Table 1 | Compositions of the

starting materials

Oxide Htn Dacite Basalt*

SiO2 68.8 2 65.58 25 50.61

TiO2 0.01 1 0.81 7 0.91

Al2O3 19.8 2 15.93 14 17.99

FeOtotal 0.00 0 5.21 25 9.13

MnO 0.00 0 1.37 5 0.16

MgO 0.00 0.19 9 6.71

CaO 3.37 8 3.99 9 11.07

Na2O 8.06 28 4.94 14 2.80

K2O 0.01 2 1.98 9 0.49

P2O5 - - 0.13

Total 100.00 100.00 100.00

Haplotonalite (Htn) and dacite glass analyses were

performed by EMPA, with standard deviation

indicated by italic font.

* XRF analysis2

6

Supplementary Table 2 | Experimental results on single compositions

Material Exp. # r

(mm) L

(mm) Total strain

Step strain

Strain rate (s

-1)

T (°C)

Jacket Jacket contrib.

Log stress (Pa)

Step viscosity

Viscosity (Pa s)

Err. max viscosity

Htn

PP300 7.46 7.21 0.4

0.09 3.9E-05 700

Cu

1% 7.8 12.08 12.08

1.1

0.02 8.9E-05 700 1% 7.8 11.86 1.0

0.06 4.6E-05 750 2% 7.3 11.67

11.56

1.0

0.15 1.0E-04 750 1% 7.6 11.56 1.0

0.10 2.1E-04 750 1% 7.8 11.46 0.9

0.01 1.3E-04 800 24% 6.3 10.20 10.20 0.7

PP231 7.49 10.31 0.9

0.07 7.0E-05 900

Fe

12% 4.6 8.76

8.37

0.5

0.11 2.0E-04 900 14% 4.8 8.45 0.4

0.08 3.5E-04 900 15% 4.9 8.31 0.4

0.15 1.3E-03 900 18% 5.1 7.98 0.4

0.02 5.1E-05 900 13% 4.3 8.62

8.17

0.5

0.06 2.1E-04 900 15% 4.5 8.14 0.4

0.06 4.3E-04 900 17% 4.7 8.05 0.4

0.08 1.3E-03 900 20% 5.0 7.85 0.4

0.09 9.8E-05 1000 33% 3.2 7.17 7.07

0.5

0.12 2.4E-04 1000 30% 3.4 6.98 0.5

0.07 7.1E-05 1100 30% - <5

Dacite PP301 7.48 8.80 0.8

0.12 4.4E-05 750

Cu

28% 5.6 9.62

9.79

0.7

0.06 8.9E-05 750 26% 5.6 9.99 0.7

0.16 1.7E-04 750 20% 6.0 9.75 0.7

0.01 8.2E-05 800 36% 5.2 9.32

9.42

0.6

0.08 9.2E-05 800 27% 5.6 9.63 0.7

0.19 1.8E-04 800 29% 5.5 9.30 0.6

0.02 8.4E-05 850 44% 4.8 8.88

8.49

0.6

0.06 3.4E-04 850 45% 4.7 8.21 0.4

0.07 7.2E-04 850 36% 5.2 8.39 0.5

Basalt

PP247 7.48 10.30 0.4

0.19 8.0E-05 1100

Fe

19% 5.7 9.83 9.70

0.6

0.13 1.6E-04 1100 19% 5.8 9.57 0.6

0.08 8.5E-05 1120 26% 4.5 8.56 8.65

0.5

0.04 3.8E-04 1120 28% 5.3 8.74 0.5

PP153 7.47 7.14 0.3

0.06 4.4E-05 1140

Fe

40% 2.0 6.38 6.14

0.6

0.06 5.1E-05 1140 42% 1.6 5.89 0.7

0.02 3.1E-05 1150 - - <5

-

0.06 1.4E-04 1150 - - -

0.09 3.1E-04 1170 - - <5 -

Exp. # refers to the number of experiment, r and L are the sample radius and length (in mm), respectively. Some experiments were conducted through different steps, with varied temperature or strain rates.

7

Supplementary Table 3 | Experimental results on mixing between dry haplotonalite and dacite compositions

Material Exp. # r

(mm) L

(mm) Total strain

Step strain

Strain rate (s

-1)

T (°C)


Log stress (Pa)

Step viscosity

Viscosity (Pa s)

Err. max viscosity

Htn + Dacite

PP302 7.40 8.17 4.3

0.04 5.9E-05 850

Fe

32% 5.0 9.23

9.11

0.6

0.06 1.3E-04 850 27% 5.3 9.15 0.6

2.77 3.3E-04 850 25% 5.5 9.01 0.5

1.41 6.4E-04 850 22% 5.9 9.05 0.5

PP303 7.48 8.20 8.3

0.05 3.5E-05 900

Fe

32% 4.1 8.59

8.23

0.5

0.15 1.0E-04 900 28% 4.4 8.36 0.5

0.95 3.3E-04 900 24% 4.6 8.04 0.4

3.46 5.4E-04 900 25% 4.7 7.94 0.4

0.11 3.3E-04 850 33% 5.3 8.80 8.82

0.5

3.53 5.4E-04 850 26% 5.6 8.84 0.5

PP349 7.47 6.46 1.7 1.70 2.1E-04 950 iron - - - 7.6 -

PP346 7.49 4.78 1.7 1.68 4.2E-04 1000 iron - - - 7.1 -

Exp. # refers to the number of experiment, r and L are the sample radius and length (in mm), respectively. Some experiments were conducted through different steps, with varied temperature or strain rates. Numbers italicized are calculated values.

Supplementary Table 4 | Experimental results on mixing between dry haplotonalite and basalt compositions

Material Exp. # r

(mm)

L (mm)

Total strain

Step strain

Strain rate (s

-1)

T (°C)


Log stress (Pa)

Step viscosity

Viscosity (Pa s)

Err. max viscosity

dry Htn+Basalt

PP149 7.47 10.48 1.9 1.93 8.9E-05 900 Fe 11% 4.7 8.76 8.76 0.5

PP235 7.47 8.73 0.5 0.50 2.9E-03 1050 Fe 40% 3.7 6.32 6.32 0.5

PP155 7.25 8.65 1.3 1.26 2.3E-04 1150 Fe - - - 4.6 -

PP156 6.91 8.64 4.9 4.92 3.1E-04 1160 Fe+steel - - - 4.4 -

PP157 6.91 8.47 3.6 3.62 4.4E-04 1170 Fe+steel - - - 4.2 -

PP160 6.91 6.31 3.1 3.10 9.3E-04 1170 Fe+steel - - - 4.2 -

PP161 6.91 5.28 3.1 3.11 4.8E-04 1170 Fe+steel - - - 4.2 -

PP164 6.91 8.27 3.9 3.93 4.7E-04 1170 Fe+steel - - - 4.2 -

PP151 7.47 9.83 1.3

0.26 7.2E-05 1200

Fe

- - -

3.5

-

0.25 2.3E-04 1200 - - - -

0.82 6.7E-04 1200 - - - -

PP176 6.92 6.64 3.8 3.78 5.2E-04 1200 Fe+steel - - - 3.5 -

Exp. # refers to the number of experiment, r and L are the sample radius and length (in mm), respectively. Some experiments were conducted through different steps, with varied temperature or strain rates. Numbers italicized are calculated values.

8

Supplementary Table 5 | Experimental results on mixing between hydrous haplotonalite and basalt compositions

Material Exp. # r

(mm) L

(mm) Total strain

Step strain

Strain rate (s

-1)

T (°C)


Log stress (Pa)

Step viscosity

Viscosity (Pa s)

Err. max viscosity

Hydrous Htn+Basalt

PP265 7.48 5.33 0.7

0.20 4.3E-05 600

copper

4% 7.3 11.64

11.42

1.0

0.21 8.5E-05 600 4% 7.3 11.32 0.9

0.33 2.3E-04 600 4% 7.7 11.31 0.9

PP258 7.46 10.93 1.7

0.03 4.2E-05 715

Cu

13% 6.5 10.85

10.77

0.8

0.14 8.8E-05 715 5% 6.6 10.69 0.8

0.85 4.8E-04 715 4% 7.5 10.84 0.8

0.69 8.1E-04 715 4% 7.6 10.69 0.7

PP285 7.41 7.39 2.4

0.02 4.3E-05 950

Fe

33% 4.3 8.69

8.57

0.5

0.33 2.9E-04 950 34% 4.9 8.44 0.4

2.03 5.9E-04 950 36% 5.0 8.22 0.4

PP295 7.35 7.37 2.0 0.69 9.6E-05 975

Fe 31% 4.2 8.26

8.14 0.4

1.34 2.0E-04 975 32% 4.3 8.03 0.4

PP296 6.89 8.21 5.1

0.45 1.0E-04 985

Fe+steel

35% 3.6 7.64

7.65

0.5

0.45 3.1E-04 985 33% 4.2 7.71 0.4

4.18 8.0E-04 985 37% 4.5 7.58 0.4

PP261 7.46 8.31 1.3

0.12 9.0E-05 1000

Fe

48% 3.1 7.48

7.17

1.0

0.06 4.3E-05 1000 49% 3.1 7.17 1.1

1.08 2.8E-04 1000 50% 3.3 6.85 1.0

PP293 6.90 8.71 0.7 0.35 2.1E-04 1020

Fe+steel 63% 2.6 6.32

6.18 1.2

0.39 8.7E-04 1020 61% 3.0 6.04 1.2

Exp. # refers to the number of experiment, r and L are the sample radius and length (in mm), respectively. Some experiments were conducted through different steps, with varied temperature or strain rates.

9

Supplementary Table 6 | Source of data for the parameters of the viscosity equation

Reference Crystal shape ratio

3 1

4 <1.5

5 1.5

6 2.4

7 3

8 4.9

9 5

Supplementary Table 7 | Symbols and parameter values used in thermal calculations

Parameter Symbol Value Unit

Final temperature T - °C

Fraction of mafic magma injected x - -

Heat capacity of felsic and mafic magmas Cf-m 1.26 J g-1

K-1

Mass fraction of crystal dissolved in felsic magma Xf variable -

Heat of fusion of crystals in felsic magma Lf 293 J g-1

Mass fraction of new crystals formed in mafic magma Xm variable -

Heat of fusion of crystals in mafic magma Lm 418 J g-1

10

Supplementary Table 8 | Liquidus temperatures and bulk water content for felsic and mafic end members, and input temperatures used for each end-

member in the mixing calculations shown in Fig. 5-7 and Supplementary Figures 3, 4.

Composition Initial T Liquidus T Bulk H2O Reference

(°C) (°C) (wt%)

0.7-1 GPa Basalt Mg 1210 1210 6.0 10

Dacite Pinatubo 750 1000 13.0 11

Dacite H2O undersat. 750 1000 9.0 11

Rhyolite HPG 750 1030 13.0 12

0.2 GPa Rhyolite HPG undersat. 750 1030 3.0 12

Greywacke 750 1060 1.8 13

Amphibolite 750 1060 4.0 14

Basalt Fe 1145 1145 6.0 1

Basalt HAB 980 980 6.0 15

Andesite Mt Pelée 890 1040 6.1 16

Dacite Pinatubo 776 930 7.0 17

Rhyolite Unzen 780 850 6.0 18

11

Supplementary References 1. Botcharnikov, R. E., Almeev, R. R., Koepke, J., & Holtz, F. Phase relations and liquid lines of descent in hydrous ferrobasalt—

implications for the Skaergaard intrusion and Columbia River flood basalts. J. Petrol. 49, 1687-1727 (2008).

2. Andújar, J., Scaillet, B., Pichavant, M., & Druitt, T. H. Differentiation conditions of a basaltic magma from Santorini and its

bearing on andesitic/dacitic magma production. In AGU Fall Meeting Abstracts (Vol. 1, p. 2354) (2010).

3. Lejeune, A. M., & Richet, P. Rheology of crystal‐ bearing silicate melts: An experimental study at high viscosities. J. Geophys.

Res. 100, 4215-4229 (1995).

4. Champallier, R., Bystricky, M., & Arbaret, L. Experimental investigation of magma rheology at 300 MPa: From pure hydrous

melt to 76 vol.% of crystals. Earth Planet. Sci. Lett. 267, 571-583 (2008).

5. Scott, T., & Kohlstedt, D. L. The effect of large melt fraction on the deformation behavior of peridotite. Earth Planet. Sci. Lett.

246, 177-187 (2006).

6. Caricchi, L., Burlini, L., Ulmer, P., Gerya, T., Vassalli, M., & Papale, P. Non-Newtonian rheology of crystal-bearing magmas and

implications for magma ascent dynamics. Earth Planet. Sci. Lett. 264, 402-419 (2007).

7. Rutter, E. H., & Neumann, D. H. K. Experimental deformation of partially molten Westerly granite under fluid‐ absent

conditions, with implications for the extraction of granitic magmas. J. Geophys. Res. 100, 15697-15715 (1995).

8. Vona, A., Romano, C., Dingwell, D. B., & Giordano, D. The rheology of crystal-bearing basaltic magmas from Stromboli and

Etna. Geochim. Cosmo. Acta, 75, 3214-3236 (2011).

9. Picard, D., Arbaret, L., Pichavant, M., Champallier, R., & Launeau, P. Rheology and microstructure of experimentally deformed

plagioclase suspensions. Geology 39, 747-750 (2011).

10. Melekhova, E., Annen, C., & Blundy, J. Compositional gaps in igneous rock suites controlled by magma system heat and water

content. Nature Geo. 6, 385-390 (2013).

11. Prouteau, G., & Scaillet, B. Experimental constraints on the origin of the 1991 Pinatubo dacite. J. Petrology 44, 2203-2241

(2003).

12. Johannes, W., & Holtz, F. Petrogenesis and experimental petrology of granitic rocks (Vol. 335). Berlin: Springer (1996).

13. Montel, J. M., & Vielzeuf, D. Partial melting of metagreywackes, Part II. Compositions of minerals and melts. Contr. Mineral.

Petrol. 128, 176-196 (1997).

14. Sisson, T. W., Ratajeski, K., Hankins, W. B., & Glazner, A. F. Voluminous granitic magmas from common basaltic sources.

Contr. Mineral. Petrol. 148, 635-661 (2005).

15. Sisson, T. W., & Grove, T. L. Experimental investigations of the role of H2O in calc-alkaline differentiation and subduction zone

magmatism. Contr. Mineral. Petrol. 113, 143-166 (1993).

16. Martel, C., Pichavant, M., Holtz, F., Scaillet, B., Bourdier, J. L., & Traineau, H. Effects of fO2 and H2O on andesite phase

relations between 2 and 4 kbar. J. Geophys. Res. 104, 29453-29470. (1999).

17. Scaillet, B., & Evans, B. W. The 15 June 1991 eruption of Mount Pinatubo. I. Phase equilibria and pre-eruption P–T–fO2–fH2O

conditions of the dacite magma. J. Petrol. 40, 381-411 (1999).

18. Tomiya, A., Takahashi, E., Furukawa, N., & Suzuki, T. Depth and evolution of a silicic magma chamber: Melting experiments on

a low-K rhyolite from Usu volcano, Japan. J. Petrol. 51, 1333-1354 (2010).

On the conditions of magma mixing and its bearing on andesite production in the crust

Documents