Decarbonization related to continental arc magmatism as a possible mechanism for Cretaceous warming

Decarbonization Related to Continental Arc

Magmatism as a Possible Mechanism for

Cretaceous Warming

by

Anna Elizabeth Brunner

MASSACHUSES INSTITUTE OF TECHNOLOGY

OCT

2 4

2017

LIBRARIES

Submitted to the Department of Earth, Atmospheric, and Planetary

Science

in partial fulfillment of the requirements for the degree of

Bachelor of Science in Earth, Atmospheric and Planetary Sciences

at the

MASSACHUSETTS INSTITUTE OF TECHNOLOGY

June 2012

@

Massachusetts Institute of Technology 2012.

Author ..

All rights reserved.

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Department of Earth, Atmospheric, and Planetary Science

Certified by...

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May 11, 2012

Accepted by..

Oliver Jagoutz

Assistant Professor, Department of EAPS

S

Thesis Supervisor

Signature redacted

...

Samuel Bowring

Chair, Committee on Undergraduate Program

Decarbonation Related to Continental Arc Magmatism as a

Possible Mechanism for Cretaceous Warming

by

Anna Elizabeth Brunner

Submitted to the Department of Earth, Atmospheric, and Planetary Sciences on May 11, 2012, in partial fulfillment of the

requirements for the degree of

Bachelor of Science in Earth, Atmospheric and Planetary Sciences

Abstract

Elevated concentrations of CO2 have been proposed as the reason that the

Creta-ceous climate was 6-14'C warmer than the present, however the source of CretaCreta-ceous CO2 is unknown [Barron, 1983]. This study examines the possibility of continental

arc magmatism as a mechanism for CO2 release, specifically as a volatile produced

during crustal assimilation and contact metamorphism of carbonates around plutons. Bedrock maps of the North American Cordillera (a region of active continental arc magmatism during the Cretaceous), the relative locations of the carbonates, the Cre-taceous plutons, and the calculated "decarbonation zones"around the plutons. These measurements were then input in a thermal and petrologoical model in order to esti-mate the quantity of CO2 released by continental arc magmatism. Testing a number

of cases with varying parameters, the model found the arc-magmatism-induced tem-perature difference between the present and Cretaceous global climates to have a lower limit of AT < 1'C and an upper limit of 5.1 < AT < 12.3'C. Decarbonation from continental arc magmatism is shown to be a possible mechanism of paleoclimatic warming, and more work is required to either confirm or refute the hypothesis. Thesis Supervisor: Oliver Jagoutz

Acknowledgments

The author would like to thank Ben Black for help preparing samples for thin section, Ben Mandler for help with thin section interpretation, Katie Pesce, Jane Connor, Jessica Ruprecht, and Simone Agha for help with revisions, and Jocelyn Fuentes for emotional support. Most of all thank you to Oli Jagoutz for being so patient with me and working with me step-by-step through such a large-scale project and also for carrying my all of my rocks in the field.

Contents

1 Introduction 13

1.1 The unsolved cause of Cretaceous warming . . . . 13 1.2 Background . . . . 14 1.2.1 Geologic context of the North American Cordillera . . . . 14 1.2.2 Metamorphic decarbonation reactions due to the formation of

calc-silicates as a mechanism for CO₂ release . . . . 18

2 Methods 21

2.1 Groundtruthing and field observations . . . . 21 2.2 Approximating the widths of decarbonation zones . . . . 23

2.2.1 Modeling the heat flow around intrusions to understand the role of contact metamorphism . . . . 23 2.3 Contact metamorphism and crustal assimilation area estimation . . . 28 2.4 Petrologic model to estimate CO₂ release temperatures and amounts 31 2.5 Extrapolating continental arc contributions of CO2 from a regional to

a global scale . . . . 33 2.6 Predicting effect on global climate . . . . 34

3 Results 37

3.1 Mineralogy of calc-silicate samples . . . . 37 3.2 _{Decarbonation zone widths . . . . 38} 3.3 Area estimation using USGS bedrock data and ArcGIS . . . . 39

4 Analysis 45

5 Discussion 49

List of Figures

1-1 Temperature throughout the Phanerozoic. . . . . 13

1-2 Distribution of Cretaceous plutons and carbonates in the North Amer-ican C ordillera. . . . . 15

1-3 Timing of formation of the major Cretaceous batholiths. . . . . 16

1-4 Phanerozoic CO2 levels. . . . . 17

1-5 Dependence of decarbonation reactions on temperature and water con-tent. ... ... ... ... ... .. 18

2-1 Mountain ranges of the Mojave Desert. . . . . 22

2-2 Diagram of an intrusion and its contact aureole. . . . . 24

2-3 Geothermal model results. . . . . 28

2-4 Thermal profile of an intrusion over time. . . . . 29

2-5 CO2 release associated with tremolite and wollastonite formation. . . 33

2-6 Orogenic belts of the Cretaceous. . . . . 34

3-1 Tremolite in Kilbeck Hills thin section. . . . . 38

3-2 Wollastonite in Kilbeck Hills thin section. . . . . 39

3-3 Garnet- in Kilbeck Hills thin section . . . . . 41

3-4 Wollastonite and garnet in Little Maria Mts thin section. . . . . 42

List of Tables

2.1 Cases considered in calcuating decarbonation zone widths . . . . 27 2.2 Representative composition of a 50% sandstone/50% limestone mixture 32 3.1 Locations of calc-silicate samples from the Kilbeck Hills and Little

M aria M ts . . . . 37 3.2 Calculated decarbonation zone widths . . . . 40 3.3 Areas for each decarbonation zone width . . . . 43 4.1 Estimated CO2 produced by global continental arc magmatism in

Cre-taceous... ... 46 4.2 Estimated global termperature changes caused by

continental-arc-magmatism-related changes in concentrations of C02, relative to present-day

Chapter 1

Introduction

1.1

The unsolved cause of Cretaceous warming

This thesis examines the hypothesis that CO2 release, or decarbonation, due to

con-tact metamorphism during the Cretaceous period made a significant contribution to the elevated temperatures of the Cretaceous climate.

10

-6

0 100 200 300 400 500 600

Tmw (Ma)

Figure 1-1: Estimated temperature of the global climate throughout the Phanero-zoic, showing the elevated temperature during the Cretaceous (adapted from

[Crowley and Berner, 2001]).

It is well-accepted that the Cretaceous climate was much more equable than that of the present-day due to fossil evidence of tropical plants at high paleolatitudes. The Cretaceous climate was warmer than the present climate by -6-14'C [Barron, 1983]. An estimate of the temperature of the global climate throughout the last 600 Ma is

shown in Figure 1-1.

There is very little agreement about what caused Cretaceous warming. Proppsed climatic forcing factors include the effect of the paleogeography on oceanic and at-mospheric circulation, increased poleward heat transport by oceanic or atat-mospheric circulation, increased atmospheric CO2 levels due to geologic processes, and even in-creased atmospheric methane levels due to megafauna [Barron and Washington, 1984, Wilkinson et al., 2012].

Geologic processes that can release atmospheric CO2 include volcanic degassing, regional metamorphism, contact metamorphism, and magmatic assimilation. Other studies have attempted to quantify the rates of CO2 release due to regional

metamor-phism; however, rates of CO2 release from continental are plutonic processes (contact

metamorphism and magmatic assimilation) have not yet been investigated.

1.2

Background

1.2.1

Geologic context of the North American Cordillera

Because Cretaceous continental arc plutons were emplaced in the same general area as Paleozoic shallow-water sedimentary rocks (especially carbonates) in the North American Cordillera and due to the availability of USGS standardized bedrock maps of the USA, the North American Cordillera is an ideal study area in which to test the hypothesis that continental arc magmatism contributed significantly to an increase in temperature in the Cretaceous climate.

Overview of the geologic history of the North American Cordillera

The North American plate was part of the supercontinent Rodinia from ~1000 to -600 Ma, until the rifting of Rodinia in the late Proterozoic to early Paleozoic cre-ated a continental shelf on the western margin of the proto-North American continent [Hoffman, 1991, Powell et al., 1993, Bond et al., 1984]. After the rifting of Rodinia this margin remained passive throughout the Paleozoic, allowing the deposition of

Distribution of Carbonates and Cretaceous Plutons

in North American Cordillera

AARB

Legend

SBB

Carbonates

Cretaceous plutons

1:35,000,000

SNB

PRB *unshaded areas in Alaska denote gaps in the USGS bedrock data

Figure 1-2: Distribution of Cretaceous plutons and carbonates in the North Amer-ican Cordillera. Bedrock geology data comes from the USGS databases, and geo-logical databases of the governments of British Columbia and the Yukon Territory [USGS, 2012, Massey et al., 2005, Gordey and Makepeace, 1999]. The bedrock geol-ogy data was incomplete in Alaska, so areas in which no data were available are left blank. The batholiths indicated are the Alaska Aleutian Range Batholith (AARB), the Coast Mountains Batholith (CMB), the Atlanta Lobe of Idaho Batholith(IB

(Atl)), Bitterroot Lobe of Idaho Batholith (IB (Bit)), the Boulder Batholith (BB),

the Sierra Nevada Batholith (SNB), and the Peninsular Ranges Batholith (PRB).

thick sequences of shallow-water sediments [Bond and Kominz, 1984]. In the Meso-zoic the dominant geologic processes on the continental margin changed drastically from passive margin subsidence and sedimentation to magmatism and terrane ac-cretion due to subduction [Burchfiel et al., 1992]. There was no east-dipping sub-duction along the western margin of the North American plate during most of the Paleozoic, so passive margin processes (continental shelf sedimentation and subsi-dence) dominated on the western margin of North America during this time period

[Burchfiel et al., 1992].

Pa-- - - -- - --- ~-- -- I

cific plate under the North American plate created a subduction-related volcanic arc in the Cordillera [Burchfiel et al., 1992]. Magmatism was limited to the narrow back-arc region of the Cordillera between the Late Permian and the Early Juras-sic (from ~250 to ~140 Ma), but the amount and the extent of magmatism in-creased during the Early Jurassic and continued through the rest of the Mesozoic [Burchfiel et al., 1992]. Possibly due to the flattening of the subducting slab during the Late Cretaceous to Early Cenozoic (from -85 to -60 Ma), the amount of arc magmatism in the western Cordillera was subdued and magmatism further inland increased [Burchfiel et al., 1992].

Major Cretaceous batholiths in the North American Cordillera

Age and Duration of Major Cordilleran

Batholiths

Cenozoic

Cretaceous

Jurassic

AARB CMB IB(Atl) IB(Bit) BB 4 SNB PRB I f - 44 40 60 80 100 120 140 160

Age (Ma)

Figure 1-3: The timing of formation of the Cretaceous major batholiths in the North American Cordillera. AARB - Alaska Aleutian Range Batholith, CMB

-Coast Mountains Batholith, IB(Atl) - Atlanta Lobe of Idaho Batholith, IB(Bit)

-Bitterroot Lobe of Idaho Batholith, BB - Boulder Batholith, SNB - Sierra Nevada Batholith, PRB -Peninsular Ranges Batholith. Ages from [Reed and Lanphere, 1973,

Gaschnig et al., 2010, Schmidt et al., 2009, Ducea, 2001].

Phanerozoic Carbon Dioxide

8000 Models Measurements ₃₀

7000 - GEOCARB I . Royer Compilation

S6000 - COPSE - 30 Myr Filter 25

5000 Rothman0 2 1 -21 4000 ₁₅ 3000-2000 NJ Pg K IJ_ Tr P C ID, IS 1 Cm I_ 0 100 200 300 400 500

Millions of Years Ago

Figure 1-4: A summary figure of the results of many different models of Phanerozoic CO2 levels, as well as measurements from CO2 proxies, illustrating the difficulty

of accurately predicting the amount of atmospheric CO2 in the distant past. The

Cretaceous (K) values of CO2, for instance, range from about 500 ppmv to more than

3000 ppmv, but qualitatively this range implies that Cretaceous atmospheric CO2

concentration was higher than that of the present-day. Adapted from [Rohde, 2005].

episodes occuring in the North American Cordillera during the Cretaceous. They are: the Peninsular Ranges Batholith (PRB), the Sierra Nevada Batholith (SNB),

the Alaska-Aleutian Range Batholith (AARB), the Idaho Batholith (IB), and the

Coast Mountains Batholith (CMB). The locations of each of these major Cretaceous batholiths, as well as that of the much smaller Boulder Batholith, are shown in Figure 1-2. These batholiths were all emplaced between the Early Cretaceous and the early Cenozoic, from -125 Ma to -40 Ma (Figure 1-3), so North American Cordilleran magmatism overlapped with the relevant period of Cretaceous warming. Based on geobarometric estimates from coexisting mineral phases, the depths of these batholiths ranges from lower bounds of 3-4 kbar for the IB [Gaschnig et al., 2010] and for some regions of the SNB [Petford et al., 2000], to upper bounds of 6-8 kbar for the CMB [McClelleland et al., 1990] and 7-9 kbar for other regions of the SNB

1.2.2

Metamorphic decarbonation reactions due to the

for-mation of calc-silicates as a mechanism for CO

release

0.2 0.4

XC

02

0.6 0.8 1.0

Figure 1-5: Phase diagram of a calc-silicate system at constant pressure (5 kbar) .

Xco2 is the fraction of the fluid phase that is C02; 1 - Xco2 is the fraction that is

H20. Adapted from [Spear, 1995].

Calc-silicate rocks are formed when carbonates and silicates react under conditions of high temperature or increased fluid flow, often releasing

CO

common reactions involved during the formation of calc-silicates is shown in Figure

1-5.

The balanced decarbonation reactions for the formation of the minerals tremolite

900

P

700

500

400

0

Tremolite

(a)

and wollastonite are [Bozhilov and Jenkins, 2001],

5(Ca,Mg)CO3 + 8SiO2 -+ 3CaCO3 + Ca2Mg5SisO2 2(OH)2 +7CO2 (1.1) dolomite quartz+H20 calcite tremolite

CaCO3 + SiO2 - CaSiO3 +C02. (1.2)

calcite quartz wollastonite

The wollastonite-forming reaction and the tremolite-forming reaction are of inter-est because petrologic modeling showed that these are the two main reactions which release CO2 in a simple system composed of a reasonable mixture of carbonate and

sandstone (see Section 2.4) and because the presence of tremolite and wollastonite in contact metamorphosed calc-silicates was verified in the Kilbeck Hills and Little Maria Mountains of the Mojave (see Section 2.1).

Chapter 2

Methods

To determine whether continental arc magmatism could have contributed signifi-cantly to Cretaceous warming, the concentration amount of CO2 released during the

Cretaceous was estimated. This estimate was made by finding the regions where decarbonation reactions could have occurred due to contact metamorphism or mag-matic assimilation of carbonates (i.e. decarbonation zones). Then the total amount of CO2 that would have been released from these regions was approximated. Finally

the regional contribution of the North American Cordillera was converted to a global amount of CO2.

To confirm the theoretical estimate of CO2 released by continental arc magmatism,

samples of calc-silicates were collected from two locations in the Mojave Desert, CA, and petrographic and field observations were made.

2.1

Groundtruthing and field observations

To support the model field observations specific to this study were made in the Kil-beck Hills and the Little Maria Mountains. Previous fieldwork has also been done by the author in two tangentially relevant areas, the Grand Canyon and the River-side Mountains (Figure 2-1). The Grand Canyon is the type location of a thick Paleozoic sedimentary sequence which contains many carbonates and sandstones, the ingredients for calc-silicates. The Kilbeck Hills, Little Maria Mountains, and

River-MIS Mf Bristol Mu Mfg MISMf =S I Mts

one "49 WIs ARIZONA

... a n STUOV AREA mm:tA1 0M4JAMLA MEXiCO 0.- 11. I I N I if Is.

I

1 k

rn-

Mfgi

plmso

merst rems MIS Wit

. W. Mxn 6 .

Figure 2-1: The mountain ranges in the Mojave Desert. Sites of field observations are highlighted in green. Calc-silicate samples were collected by the author from the Kil-beck Hills and Little Maria Mountains. Adapted from [Spencer and Reynolds, 1990].

side Mountains each have an analogous stratigraphy to the Grand Canyon Paleozoic sequence; however, the units are deformed and metamorphosed. Calc-silicate miner-als have been identified in the Kilbeck Hills and the Little Maria Mountains in the metamorphosed Paleozoic sequence. The Grand Canyon Paleozoic sequence in the Riverside Mountains, on the other hand, has been metamorphosed and highly de-formed, but calc-silicates were observed so no decarbonation reactions have occurred. In the Kilbeck Hills and the Little Maria Mountains, the minerals wollastonite and tremolite (which indicate decarbonation) were identified in the field. Samples were collected to confirm the presence of these minerals and assess whether the assumption that calc-silicates were created by Cretaceous contact metamorphism is reasonable.

Six calc-silicate samples were made into thin sections, four from the Kilbeck Hills and two from the Little Maria Mountains (Table 3.1). The 30-pm-thick thin sections of the samples were examined under polarized microscopes, in order to identify the minerals present.

2.2

Approximating the widths of decarbonation

zones

To find the extent of decarbonation reactions due to contact metamorphism within the Cretaceous continental arc of the North American Cordillera, an estimate for the decarbonation zone widths was made using a thermal model around an intrusion. The approximated decarbonation widths were then used to determine the quantity of carbonates within that distance of a Cretaceous intrusion.

2.2.1

Modeling the heat flow around intrusions to

under-stand the role of contact metamorphism

In regions undergoing continental arc magmatism, nearby rocks are contact meta-morphosed due to increased temperatures and sometimes due to increased fluid flow. Depending on the initial mineralogy of the host rocks contact metamorphism may produce

CO

metamorphism, and thus constrain the amount of CO2 produced.

The contact aureole for an intrusion is the zone within the host rock which ex-periences metamorphism due to the heat of the intrusion. If the host rocks contain carbonates, decarbonation reactions may occur. The occurrence of these reactions depends upon conditions including temperature, so the higher-temperature wollas-tonite reaction will only occur closer to the intrusion where temperatures are higher; whereas, the lower-temperature tremolite reaction occurs further from the intrusion. The widths of the tremolite and wollastonite zones (or decarbonation zones) in the contact aureole represent the distances away from the intrusion where the surrounding

twPth

ST

Mdh

Figure 2-2: The contact aureole for an intrusion is the zone within the host rock which experiences metamorphism due to the heat of the intrusion. If the host rocks contain carbonates, decarbonation reactions may occur. The occurrence of these reactions depends upon conditions including temperature, so the higher-temperature wollastonite reaction will only occur closer to the intrusion where temperatures are higher; whereas, the lower-temperature tremolite reaction occurs further from the intrusion. Lithologic patterns used are from [FDGC, 2006].

N1

th

No' '

rock reached the temperatures necessary for contact metamorphism (Figure 2-2). The decarbonation zone width depends on the temperature of the intrusion T, width of the intrusion wist, temperature of the surrounding rock Tb, and the temperature at which decarbonation reactions proceed Tco2. T is a function of the depth of the

intrusion zit and the geotherm (temperature as a function of depth). Tco2 depends

on the presence of fluids, the water content of those fluids XH20, and the specific

mineral reaction involved.

To model the heat transport around an intrusion, a one-dimensional intrusion is assumed and temperature was calculated as a function of horizontal distance away from the intrusion (intrusion width, wint). To simplify the problem only heat con-duction was considered as a possible transport mechanism. In reality the circulation of fluids is likely responsible for a nontrivial portion of the heat flow surrounding the intrusions.

For temperature T, time t, thermal diffusivity K, position r, internal heat pro-duction A, density p, specific heat c, and uplift rate of U,, the full heat equation is,

dT

_T

_A

=T+ - - U O (2.1)

dt aZ2 PC ((r'

conduction heat production advection

[Spear, 1995] This equation was simplified by assuming conduction is the only method of heat transfer (i.e., that A and U, are zero) [Spear, 1995]. It then reduced to,

dT _d 2T

=T

__' . _(2.2)

dt dx2

The solution to which is [Spear, 1995]:

1 _____ -(xwint +xN

T(x) = Tb + - (T - Tb) erf + erf i (2.3)

2 2 V-, 2

v-K-with , = 32 km2_{/Ma [Stuwe, 2002].}

The T can be determined from a model of a geotherm. Two regions were consid-ered: the upper crustal region (0 km < z < 15 km) which was assumed to have an

exponential distribution of internal heat production, and the lower crustal region (15 km < z < 35 km) which was assumed to have a constant internal heat production. The geotherm is assumed to be steady-state (dT/dt = 0) and that there is no uplift

(Uz = 0), so Equation 2.1 reduced to,

2_T

= A

(2.4) OZ2 k'

since r, = k/(pc) [Spear, 1995]. The equation describing the geotherm in the upper crust is

T(z) = zq, AOD2(1 _ e-z/DA), (2.5)

k k

where AO is the internal heat production at the surface (Ao = 2x10- 6 W/m3), q, is surface heat flow, z is depth, k is thermal conductivity (k = 3.0 w), and DA is the depth constant for heat production (DA = 15 km) [Chapman, 1986, Spear, 1995]. The solution for the lower crustal geotherm is,

Az2 _(qAD\

T(z) 2 + z - + ) (2.6)

2k (k k

where A is the internal heat generation (A = 4.5 x 10-5 W/m3), q, is surface heat flow, z is depth, k is thermal conductivity (k = 3.0 W), and D is the thickness of the layer (D = 35 km) [Chapman, 1986, Spear, 1995].

For each factor which the decarbonation zone width ultimately depends on (q.,

Zint, Tint, XH2O, and Tco2), a minimum, a reasonable intermediate, and a

maxi-mum condition were selected, and 13 cases were generated using different combina-tions of condicombina-tions. Table 2.1 shows the assumed condicombina-tions used in each of the 13 cases. Case 1 uses all the minimum conditions, Case 2 all the intermediate condi-tions, and Case 3 all the maximum conditions. For the other cases, one variable was changed and all other variables were held at their intermediate values. Sur-face heat flow values of 50, 70, and 90 mW/m2

were used. The minimum q, of 50 mW/M2

is on the order of the heat flow of relatively old continental crust, the mid-dle value of 70 mW/m2

Table 2.1: Cases considered in calcuating decarbonation zone widths

Case qs zint T Tint XH2O Tco, ('C) wint

(mW/m2_{) (km) (}0_{C) (}0_{C) % tr wo (km)} 1 50 11.5 160 780 0 640 940 0.05 2 70 19 370 890 80 510 760 3 3 90 28.5 720 1000 100 400 500 20 4 50 19 230 890 80 510 760 3 5 90 19 500 890 80 510 760 3 6 70 19 370 780 0 510 760 3 7 70 19 370 1000 100 510 760 3 8 70 11.5 240 890 80 510 760 3 9 70 28.5 530 890 80 510 760 3 10 70 19 370 890 80 640 940 3 11 70 19 370 890 80 400 500 3 12 70 19 370 890 80 510 760 0.05 13 70 19 370 890 80 510 760 20

and the maximum of 90 mW/m2

is on the order of the heat flow near a mid-ocean ridge [Chapman, 1986, Davies and Davies, 2010]. The intrusion temperatures used were the reasonable minimum and maximum temperatures for granitic intrusions [Petford et al., 2000]. The depths used were chosen based on geobarometry of the major Cordilleran batholiths (see Section 1.2.1). The phase diagram from which the decarbonation reaction temperatures were selected is shown in Figure 1-5.

The calculated geotherml from which T was extracted is shown in Figure 2-3. Using Equation 2.3, the decarbonation zone widths for both of the decarbonation reactions over a reasonable range of intrusion temperatures, intrusion widths, intru-sion depths, surface heat flows, and water contents of the melt.2

Figure 2-4 is an example of the intrusion model, using the conditions for Case 2 (see Table 2.1).

Three potentially important factors which were not taken into account in the decarbonation zone model are, 1) heat released by the decarbonation reactions and, 2) heat of fusion from the crystallization of magma [Spear, 1995], and 3) heat transported

'This geotherm was calculated using a Matlab script (called "steady state geotherms with heat production") which generates a geotherm for a specified surface heat flow and intrinsic heat produc-tion rate [Reiners and Ehlers, 2005].

2

The thermal intrusion model was calculated using a Matlab script called "intrusion iD thermal response" [Reiners and Ehlers, 2005].

0

10

30

Steady-State Geotherm

0

2

3

4

5

6

7

8

9

...

.

0

100 200 300 400 500 600 700 800 900 1000

Temperature

Figure 2-3: The steady-state geotherm was calculated for three surface heat values.

flow (q,)

away from the intrusion by fluids flowing through the rock. These three factors would increase the temperatures of the host rock and thus the width of the decarbonation zones.

2.3

Contact metamorphism and crustal

assimila-tion area estimaassimila-tion

In order to estimate the volume of CO2 released by continental arc magmatism, I

found the area where the processes of contact metamorphism and magmatic assim-ilation could have occurred. Assimassim-ilation could have occurred at the edges of what are now Cretaceous intrusions. Contact metamorphism could have occurred in the overlap between pre-Cretaceous carbonate rocks and decarbonation zones around the Cretaceous intrusions.

The spatial distribution of the relevant units was mapped using the USGS bedrock

900

_-TiII-0

yars

Thermal

Profile Over

Time for Case 2

--- 1000

800

_-.-

_10,000

100,000

6700 -

1

Ma

conditions:

S600-...

19 km

T

500 -T2)

.

.w1.2

T

890

C

400

'K

Tb

300

1

I

0

1

2

3

4

5

Distance from Center of Intrusion (km)

Figure 2-4: Example demonstrating how the decarbonation zone width was deter-mined for the initial conditions of Case 2. The initial temperature profile was cal-culated at various time steps based on the solution to the heat equation assuming conduction was the only method of heat transport.

maps of Arizona, California, Idaho, Nevada, Utah, Oregon, Montana, Washington, and Alaska [USGS, 2012], and non-USGS bedrock maps of British Columbia, and the Yukon Territory [Massey et al., 2005, Gordey and Makepeace, 1999].

The age constraint of 65-144 Ma (i.e. Cretaceous period) for the igneous intru-sions means the theoretical duration of CO2 release is 80 Ma. Because the USGS

and Canadian maps usually did not specify a more constrained age range for the mapped units, a shorter duration of CO2 release is not within the scope of this

study, but might be more realistic. For instance, most of the volume of the Sierra Nevada Batholith was emplaced within 20 Ma [Ducea, 2001]. Based on the model of [Kerrick and Caldeira, 1993], a shorter duration of CO2 release would create a

stronger effect on the climate given the same quantity of released CO2 [Kerrick and Caldeira, 1993]. All the blue areas mapped in Figure 1-2 had either a primary or secondary rock

type of limestone, dolostone, or marble. The reason the presence of carbonates in

the units was required but not necessarily silicate components is that the CO2 ulti-mately comes from the carbonates, whereas silica could have been introduced by fluid exchange with the magma.

All files were projected using an Albers Equal Area Conic Projection modified to reduce the distortion for western North America, the region of interest. Since the projection is an equal area projection, area is preserved, so the calculation of area will be accurate.

Within ArcMap, the relevant units were determined using the Select By Attributes tool. Within the USGS datasets the units were selected by their UNITAGE, ROCK-TYPE1, and ROCKTYPE2 attributes. First, I created a set of polygons made out of the set of units with a carbonate rock type and an age older than Cretaceous. Next, I created a layer for the Cretaceous plutons, which contains all the rock units with a plutonic rock type and a Cretaceous age.

The Buffer tool in ArcMap's Geoprocessing menu extended the polygons in the Cretaceous pluton layer out to the average/representative distance of contact meta-morphism. The overlap between any of the buffered polygons was eliminated by using the Dissolve tool in the Geoprocessing menu to turn the entire set of polygons into one polygon.

The Intersect tool calculated the overlap between the buffered-and-dissolved Cre-taceous pluton polygons and the Paleozoic sedimentary polygons.

To convert these areas to volumes, an average sedimentary sequence thickness of 1 kilometer was assumed. For representative intrusion widths (wiwt), a minimum of 0.5 kilometers, an average of 3 kilometers, and a maximum of 20 kilometers were chosen based on measurements from ArcMap.

2.4

Petrologic model to estimate CO

release

tem-peratures and amounts

The volume found with ArcMap must be converted to a volume of CO2. To find this

volume, a representative composition must be selected, and a petrological model of the system must be constructed to see how a rock of such composition would behave in the pressure and temperature conditions of contact metamorphism. The system's behavior might also vary with water content, which was examined as well.

Since the formation of calc-silicates requires silica and CO₂ the representative composition should contain a mixture of silicates and carbonates. Both of these types of rock form significant parts of the Paleozoic sedimentary sequence, often found stratigraphically adjacent to each other (in the Grand Canyon sequence, the Kaibab Limestone and Coconino Sandstone are adjacent, as are the Muav Limestone and Tapeats Sandstone), and the units themselves can be a mixture of sandy material and carbonate material (in the Grand Canyon sequence, the Supai group is alternating layers of sandstone and sandy limestone, and the Kaibab Limestone contains silica-rich chert nodules).

Whole-rock analyses from the literature [Cox and Lowe, 1996, Ingamells and Suhr, 1967] were averaged to obtain reasonable compositions for sandstone and limestone. The limestone sample studied by [Cox and Lowe, 1996] was selected as the representative limestone composition in spite of its atypically pure composition, because it was the best complete report of whole-rock analysis of a carbonate that could be found in the literature. A 50:50 mixture by weight of the averaged sandstone composition and av-eraged limestone composition is considered the representative composition Table 2.2. To figure out what reactions occur for the simplified system presented in Table 2.2, the metamorphic system was modeled with Perplei( (version 6.6.6). PerpleX is a suite of programs created by James Connolly for modeling mineralogical systems un-der various temperature and pressure conditions in orun-der to generate phase diagrams

[Connolly, 2005].

observa-Table 2.2: Representative composition of a 50% sandstone/50% limestone mixture component weight % CO2 22.46 CaO 20.94 MgO 5.66 SiO2 47.90 A1203 1.74 FeO 0.83 K20 0.42 Na20 0.06 H20 0.00

tions at the Kilbeck Hills (see Sections 2.1 and 3). Presence of the mineral garnet in some of the intrusions and skarns in the Kilbeck Hills implies that the rocks were at a pressure of at least 3-4 kbar , so PerpleX calculations were done at 3 kbar to simulate reasonable metamorphic conditions.

Based on the PerpleX calculations using the representative composition and rea-sonable metamorphic conditions, two decarbonation reactions were found to release

C02: the tremolite-producing reaction (Eq. 1.1) and the wollastonite-producing

re-action (Eq. 1.2). As shown in Figure 2-5, the tremolite rere-action releases 10 wt% CO₂ (which is half of the 20 wt% CO2 assumed in the representative composition), and the wollastonite reaction releases the remaining 10 wt% CO₂.

The temperature at which the calc-silicate reactions occur is highly dependent on water content of the fluid phase. Water could be introduced from the subduction zone magma. Based on Figure 1-5 from [Spear, 1995], which was calculated at 5 kbar, the reaction temperatures are 640 C for the tremolite reaction and 940 C for wollastonite reaction when the water content is 0% H20 (minimum, xCO2 is 1), 510'C and 760'C

when the fluid phase is 80% H20 (XCO2 is 0.2), and 400'C and 500

0

C when the fluid phase is 100% H20 (maximum, xCO2 is 0).

40 W - - --- I I

CO2(wt%) released in calc-silicate decarbonization reactions

20

0 4000 P(bar)

20 4

0 800 1000 1200 2000

Figure 2-5: The tremolite reaction releases 10 wt% C02, and the wollastonite reaction

releases the remaining 10 wt% CO2. Tremolite forms at a lower temperature than

wollastonite.

2.5

Extrapolating continental arc contributions of

CO

from a regional to a global scale

Cretaceous continental arc magmatism was not limited to the North American Cordillera, as discussed in Section 1.2.1. In order to produce an estimate more applicable to the global atmosphere, the volume of CO2 from western North America can be scaled by

the ratio of the length of the North American continental arc to the length of the Circum-Pacific orogenic belt. The Cretaceous continental arc system consisted of the Circum-Pacific orogenic belt and the Tethyan orogenic belt [Dickinson, 2004]. The Tethyan continental arc ran from Taiwan to Europe along what is now the Alpine-Himalayan collision belt, and the Pacific continental arc extended almost all the way around the Pacific from Antarctica to Taiwan. The combined length of the Pacific and Tethyan belts was approximately 34,000 km long (Figure 2-6). The North American continental arc is approximately 5400 km long, thus the ratio is ~6.3.

[Antrctics not shown] Midagsoa_

GOO WA NA

Fir 2indg te ofheAre-Hiaesen [ii on 04

Indone2s0

.6redictingef

t n g lt

Figure 2-6: Schematic showing the orogenic belts of the Cretaceous. The Tethyan arc was ~9000 km long, and the Pacific arc was ~25,000 km long. Figure from

[Dickinson, 2004].

2.6

Predicting effect on global climate

To determine the increase in paleoclimatic temperature due to the release CO2 from

continental arc magmatism, the estimated area of continental-arc magmatism was converted into a volume of CO2 (Vco2) which in turn was converted to an atmospheric

concentration of CO2 (PCO2(arc)).

The mass of released CO2 is calculated as follows:

Mco

=

where a 0o is the area of the wollastonite decarbonation zone of the contact aureole,

atr is the area of the tremolite decarbonation zone of the contact aureole, dsed is the thickness of carbonate-containing sedimentary rock, and pc is the average density of the crust. The wollastonite decarbonation zone area awo was multiplied by a factor of 0.2 because 20% by weight is the amount of CO2 released by the wollastonite

reaction for the assumed rock composition (see Section 2.4 and Figure 2-5); similarly, the tremolite decarbonation zone area at, was multiplied by 0.1 because 10 wt% CO2

the assumed composition (Table 2.2) at the host rock temperature T and depth zint conditions specified for the most average case (Case 2, see Table 2.1), based on a Perple_X calculation.

This mass was converted to a volume (in m3) by dividing by the density of gaseous C02:

VCo2= 6.3 MC02 (2.8)

PCO2(gas)

PCO2(gas) is 238.7 kg/m3, at 25'C and standard pressure [Hodgman, 1962]. The factor

of 6.3 comes from the ratio of the North American Cordillera arc length to the global arc length (as discussed in Section 2.5).

To convert Vco2 to an atmospheric concentration (in ppmv), a reasonable volume

of atmosphere must be approximated. The size of the atmosphere is not well-defined , so I have assumed that the height of atmosphere is the atmospheric scale height (the altitude at which pressure has dropped off by a factor of e). The scale height of Earth's atmosphere hscae is 8.5 km and the average radius of the Earth REarth is 6378 km [Williams, 2010]. Therefore the volume of the atmosphere is,

4

Vatm = -7r((REarth + hscale )3 - (hscale)3) = 4.4 x 109km3 = 4.4 x 102'L, (2.9) 3

and the concentration of CO₂ (in ppmv) is,

pCO2 = VC02 x 106. (2.10)

Vatm

The present-day CO2 concentration was defined to be PCO2(original), since the

aim is to quantify the temperature change relative to the present-day climate. The Cretaceous CO2 concentration PCO2(new) was defined to be

PCO2(new) = PCO2(origina) + PCO2(arc). (2.11)

Using the climate sensitivity factor AT2. (the temperature difference if the

concen-trations (i.e. PCO2(new) and PCO2(original)), the change in the climate's temperature AT resulting from the difference between those two concentrations is described by [Archer and Buffett, 2005]:

AT = AT2, In PC0 2(new) x 1n2, (2.12)

PC0 2(original)

where 2.5 < AT22 < 4C [Hegerl et al., 2007].

With the calculated ATs, the theoretical change in Cretaceous global climate due to continental arc magmatism can be evaluated to determine whether it matches with the actual temperature of the Cretaceous climate.

Chapter 3

Results

3.1

Mineralogy of calc-silicate samples

In the field, many calc-silicate rocks were observed. Sample containing tremolite and calcite were found in the Kilbeck Hills (ABLM2, ABKH8 Figure 3-1)-the products of the tremolite reaction besides CO2. Samples containing wollastonite, calcite, and

quartz-the products and reactants of the wollastonite reaction besides CO2-were

found in both the Kilbeck Hills (ABKH18, ABKH21, Figure 3-2) and the Little Maria Mountains (ABLM1 and ABLM2, Figure 3-4). Besides the microscopic evidence of decarbonation reactions in calc-silicate rocks, there was also macroscopic evidence

(Figure 3-5).

Samples ABLM1, ABLM2, ABKH18, and ABKH21 are interpreted to have been Table 3.1: Locations of calc-silicate samples from the Kilbeck Hills and Little Maria Mts

sample name latitude longitude

ABLM1 3354'10.26"N 11459'7.04"W ABLM2 3354'10.26"N 11459'7.04"W ABKH8 3420'45.86" N 11518'29.66"W ABKH9 3420'46.74" N 11518'32.47"W ABKH18 3421'28.30"N 11517'32.57"W ABKH21 3421'28.30"N 11517'32.57"W

Figure 3-1: Tremolite (tr) and calcite (cc) in sample ABKH8 from the Kilbeck Hills.

part of the Supai Group unit prior to metamorphism. Supai consists of continuous layers of sandstone and limestone, so silica and carbonates are both present. ABKH9 is a skarn related to a nearby gabbroic intrusion, so fluid exchange definitely was responsible for some of the calc-silicate formation in the Kilbeck Hills, but certainly not all. ABKH8 is interpreted to have been part of the Temple Butte Dolomite unit prior to metamorphism. Temple Butte consists of dolomite with discontinuous chert nodules [ref].

In thin section, epidote grains were often found as inclusions within larger wollas-tonite grains.

3.2

Decarbonation zone widths

The widths of the decarbonation zones were calculated for 13 different cases, as explained in Section 2. Each case has two decarbonation zone widths, one per mineral decarbonation reaction, because the two reactions proceed at different temperatures

(Figure 2-5).

The calculated decarbonation zone widths are reported in Table 3.3. Cases 3 (maximum for all conditions) and 9 (maximum intrusion depth) are excluded because

Figure 3-2: Wollastonite (wo) and calcite (cc) in ABKH21 from the Kilbeck Hills.

the background temperature at the depth of the intrusion was greater than the mineral reaction temperatures, so the decarbonation zone width would be infinite, hence I did not continue these calculations after figuring out that the decarbonation zone did not have a meaningful definition in these cases. Cases 1 (minimum for all conditions), 10 (minimum water content), and 12 (minimum intrusion width) were all cases where the width of the decarbonation zone was zero for both mineral reactions, and only case 11 (maximum water content) had a non-zero decarbonation zone width for wollastonite.

3.3

Area estimation using USGS bedrock data and

ArcGIS

The plutons shown in Figure 1-2 were expanded by each of the decarbonation widths listed in Table 3.3, and the resulting areas were calculated in ArcMap. In order to account for the maximum effect of of total magmatic assimilation, the "total plu-ton"values are derived by assuming that 100% of the plutons intruded directly into carbonates, and that there was 100% assimilation of these host carbonates within the area of the pluton in addition to CO2 release by contact metamorphism within the

Table 3.2: Calculated decarbonation zone widths case mineral reaction Wdecarb (km)

1 trem 0 wo 0 2 trem 1.20 wo 0 3 trem x wo x 4 trem 0.23 wo 0 5 trem 4.89 wo 0 6 trem 1.20 wo 0 7 trem 0.63 wo 0 8 trem 1.77 wo 0 9 trem x wo 0 10 trem 0 wo 0 11 trem 5.35 wo 1.41 12 trem 0 wo 0 13 trem 2.20 wo 0

Figure 3-3: Large amounts of garnet (gt) and epidote (ep), and minor amounts of calcite in ABKH9, a skarn sample from the Kilbeck Hills.

aureole" results assume that all of the decarbonation zone induced contact metamor-phism of host carbonate rocks, but that no carbonates were assimilated within the pluton area. The "intersection" results only consider the area of intersection between the decarbonation zones and the carbonate rocks. The intersection approximation is the lower bound.

In reality, both assimilation of host carbonates and contact metamorphism are active processes around intrusions which are near host carbonates, though not nec-essarily at such extremes as were assumed here. This case study aimed to estimate values for the minimum and maximum end-members, as well as a number of in-between possibilities, but did not seek to uncover a single "most reasonable" estimate of how much CO2 was released in the Cretaceous.

Figure 3-4: Large wollastonite (wo) grains, small garnet (gt) and epidote (ep) grains, and interstitial quartz (qz) in ABLM1 in the Little Maria Mountains. This sample seemed fairly homogeneous at the hand sample scale.

I" i.~At

Figure 3-5: Reaction rim of wollastonite at the boundary between quartz and calcite (hammer for scale), indicating release of CO2. This unit was interpreted to be a

metamorphosed part of the Supai group in the Kilbeck Hills.

Table 3.3: Areas for each decarbonation zone width case total pluton (km2

) total aureole (km2 ) intersection (km2 ) tr wo tr wo tr wo 1 241635 241635 0 0 0 0 2 438366 241635 196731 0 11007 0 4 317385 241635 75750 0 2874 0 5 786553 241635 544918 0 37522 0 6 372640 241635 131005 0 6357 0 7 498595 241635 256961 0 15507 0 8 323172 241635 81537 0 3214 0 10 241635 241635 0 0 0 0 11 824835 460760 583200 219125 40406 12664 12 241635 241635 0 0 0 0 13 541965 241635 300331 0 18814 0

Chapter 4

Analysis

The total pluton, total aureole, and intersection results are reported in Table 4.1 as both a mass of CO2 and as a concentration of CO2 in volumetric ppm (ppmv), based

on calculations explained in Chapter 2.

In order to better understand the effect of this concentration of CO2 on the global

climate, Eq. 2.12 was used to generate approximate temperature changes based on the change in pCO2 due to continental arc magmatism as reported in 4.1. Estimated

Table 4.1: Estimated CO2 produced by global continental arc magmatism in Cretaceous

MC0 2 (kg) pCO2 (ppmv)

case total pluton total aureole intersection total pluton total aureole intersection

1 1.97652E+17 0 0 1200 0 0

2 2.51293E+17 5.36408E+16 3.00E+15 1525 325.6 18.2

4 2.18306E+17 2.06541E+16 7.84E+14 1325 125.4 4.8

5 3.4623E+17 1.48577E+17 1.02E+16 2101 901.7 62.1

6 2.33372E+17 3.57199E+16 1.73E+15 1416 216.8 10.5

7 2.67715E+17 7.00629E+16 4.23E+15 1625 425.2 25.7

8 2.20E+17 2.2232E+16 8.76E+14 1335 134.9 5.3

10 1.97652E+17 0 0 1200 0 0

11 4.76161E+17 2.78509E+17 1.45E+16 2890 1690.3 87.8

12 1.97652E+17 0 0 1200 0 0

Table 4.2: Estimated global temperature changes caused by continental-arc-magmatism-related changes in concentrations of CO2, relative to present-day

atmo-sphere

case total pluton AT ('C) total aureole AT (0_{C) intersection AT (}0

C) min max min max min max

1 5.07 8.11 0 0 0 0 2 5.74 9.18 2.19 3.50 0.16 0.26 4 5.34 8.55 1.01 1.61 0.04 0.07 5 6.69 10.70 4.32 6.91 0.53 0.85 6 5.53 8.85 1.59 2.55 0.10 0.15 7 5.92 9.48 2.66 4.25 0.23 0.37 8 5.36 8.58 1.07 1.71 0.05 0.08 10 5.07 8.11 0 0 0 0 11 7.68 12.29 6.04 9.66 0.73 1.17 12 5.07 8.11 0 0 0 0 13 6.05 9.68 2.96 4.74 0.28 0.44

Chapter 5

Discussion

The total pluton area assumption results in ATs between 5 to 12.3'C for all calculated cases. The maximum concentration of CO2 from continental arc magmatism is 2890

ppm (for case 11, total pluton), which would result in a temperature rise of 7.7-12.3'C. The Cretaceous. was -6-14'C warmer than the present [Barron, 1983], so this result suggests that continental arc magmatism may have been able to release enough CO2 to cause Cretaceous warming. However, the total pluton area estimation is unreasonable because it includes the assumptions that all of the plutons intruded directly into a 1-km-thick layer of carbonates, and that the entire volume of the carbonate host rock was assimilated (and therefore that all of the CO2 it contained was released). In addition, case 11 was calculated assuming that there was a saturated fluid phase containing 100% water. Whil some amount of water is expected because arc magmatism is usually generated when water from the subducting plate lowers the melting temperature of the overlying mantle, expecting the system to be saturated is overly idealized and even if saturation did occur in some areas it would not apply across such a large expanse as the North American Cordillera.

The intersection area assumption results in ATs which are between 0.04 to 1.2'C for all nonzero cases. This is the lower limit because it only included areas where calculated decarbonation zones around plutons overlapped with carbonates. The intersection area assumption most likely underestimates the amount of CO2 for a

car-bonates, the area estimation accounted for the contact metamorphism around the edges of the pluton but not around roof of the pluton, and the decarbonation zone width calculation did not account for heat transport away from the pluton by fluid flow which is likely to have been an important factor in some areas. Fluid flow as a heat transport mechanism around intrusions would have simulataneously increased distance away from intrusions which reached a certain temperature and the water in the fluids would have lowered the temperature required for the reaction to proceed.

Case 2 is considered to be the average case, because the defined average or rea-sonable intermediate values were used for all the inputs, and the intersection result is considered the most conservative of the three area estimates, because only the area of intersection between the decarbonation zones and the carbonates was included. The most conservative result for case 2 (intersection) is 18.2 ppmv, which corresponds to a temperature increase of 0.2-0.3'C, and the least conservative result for case 2 (total pluton) is 1525 ppmv causing a temperature change of between 5.7-9.20C.

The nonzero minimum is 4.8 ppmv, or a AT of 0.04-0.07"C (case 4). Case 4 is the case of minimum surface heat flow (50 mW/m2_{). The average continental}

surface heat flow as reported by [Spear, 1995] is 56.5 mW/M2_{, so a surface heat flow}

such as the chosen minimum is perfectly reasonable for an old continental shield, but younger continental crust would have a higher heat flow [Davies and Davies, 2010]. Additionally, 70.9 mW/m2 _{is the preferred average continental heat flow given by}

[Davies and Davies, 2010], which means this value is definitely unrealistic, but also sets a reasonable lower bound.

Chapter 6

Conclusions

The presented model indicates that decarbonation due to continental arc magmatism is a possible cause of Cretaceous warming, and certainly needs to be investigated further. The lower limit for AT is 0.04 to 1'C, which would not have caused a significant effect on the Cretaceous climate, but the upper limit for AT is 5 to 12'C, which is well within the range of estimated global temperatures in the Cretaceous.

This study shows that amount of CO2 depends primarily on how the area of

de-carbonation is calculated. It is unclear which area estimate is most accurate, but the intersection area estimate most likely underestimates the actual area of decarbona-tion, and the total pluton area estimate is most likely an overestimate. Because all the area calculations were done in map view, the plutons, carbonates, and decarbon-ation zones were all considered two-dimensional quantities, until the calculated area was multiplied by the assumed average sediment thickness ded. Therefore this very simplistic model did not account for the decarbonation from contact metamorphism and/or crustal assimilation at the roof of the pluton. By making a model for the average 3D shape of a pluton and extrapolating that shape vertically for each plu-ton, one could determine where the plutons and their decarbonation zones existed in three-dimensional space. However, figuring out the location of the carbonate layers relative to the plutons would not be a trivial task in three dimensions because the overall attitude of each of the carbonate layers would need to be known.

Bibliography

[Archer and Buffett, 2005] Archer, D. and Buffett, B. (2005). Time-dependent re-sponse of the global ocean clathrate reservoir to climatic and anthropogenic forcing.

Geochemistry Geophysics Geosystems, 6(3).

[Barron, 1983] Barron, E. J. (1983). A warm, equable cretaceous: the nature of the problem. Earth-Science Reviews, 19:305-338.

[Barron and Washington, 1984] Barron, E. J. and Washington, W. M. (1984). The role of geographic variables in explaining paleoclimates' results from cretaceous climate model sensitivity studies. Journal of Geophysical Research, 89:1267-1279. [Bond and Kominz, 1984] Bond, G. C. and Kominz, M. (1984). Construction of

tec-tonic subsidence curves for the early paleozoic miogeocline, southern canadian rocky mountains: Implications for subsidence mechanisms, age of breakup, and crustal thinning. GSA Bulletin, 95:155-173.

[Bond et al., 1984] Bond, G. C., Nickeson, P. A., and Kominz, M. A. (1984). Breakup

of a supercontinent between 625 ma and 555 ma: new evidence and implications for continental histories. Earth and Planetary Science Letters, 70:325-345.

[Bozhilov and Jenkins, 2001] Bozhilov, K. N. and Jenkins, D. M. (2001).

Analyti-cal electron microscopy of tremolite. Modern Research and Educational Topics in

Microscopy.

[Burchfiel et al., 1992] Burchfiel, B. C., Cowan, D. S., and Davis, G. A. (1992). The

Geology of North America, volume G-3, The Cordilleran Orogen: Conterminous

U.S., chapter Chapter 8: Tectonic overview of the Cordilleran orogen in the western United States. Geological Society of America.

[Chapman, 1986] Chapman, D. S. (1986). Thermal gradients in the continental crust.

Geological Society, London, Special Publications, 24:63-70.

[Connolly, 2005] Connolly, J. A. D. (2005). Computation of phase equilibria by linear programming: A tool for geodynamic modeling and its application to subduction zone decarbonation. Earth and Planetary Science Letters, 236:524-541.